Encyclopedia of soil science, Volume 1

Acid Mine Drainage Jerry Bigham Wendy Gagliano The Ohio State University, Columbus, Ohio, U.S.A. INTRODUCTION MINE DR...

Author: Rattan Lal

1675 downloads 3217 Views 43MB Size Report

This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!
Report copyright / DMCA form

DOWNLOAD PDF

Acid Mine Drainage Jerry Bigham Wendy Gagliano The Ohio State University, Columbus, Ohio, U.S.A.

INTRODUCTION

MINE DRAINAGE CHEMISTRY

What Is Acid Mine Drainage?

Mine drainage is a complex biogeochemical process involving oxidation-reduction, hydrolysis, precipitation, and dissolution reactions as well as microbial catalysis.[1] The entire sequence is commonly represented by Reaction (1), which describes the overall oxidation of pyrite by oxygen in the presence of water to form iron hydroxide [Fe(OH)3] and sulfuric acid.

Acid mine drainage refers to metal-rich sulfuric acid solutions released from mine tunnels, open pits, and waste rock piles (Table 1). Similar solutions are produced by the drainage of some coastal wetlands, resulting in the formation of acid sulfate soils. Acid mine drainage typically ranges in pH from 2 to 4; however, extreme sites like Iron Mountain, California have produced pH values as low as 3.6.[1] Neutral to alkaline mine drainage is also common in areas where the surrounding geologic units contain carbonate rocks to buffer acidity (Table 1). Why Is Acid Mine Drainage a Problem? Soils and spoils exposed to acid mine drainage do not support vegetation and are susceptible to erosion. When acid mine drainage enters natural waterways, changes in pH and the formation of voluminous precipitates of metal hydroxides can devastate fish populations and other aquatic life (Fig. 1). The corrosion of engineered structures like bridges is also greatly accelerated. There may be as many as 500,000 inactive or abandoned mines in the United States, with mine drainage severely impacting approximately 19,300 km of streams and more than 72,000 ha of lakes and reservoirs.[2,3] Once initiated, mine drainage may persist for decades, making it a challenging problem to solve.

3 1 FeS2ðsÞ þ 3 O2ðgÞ þ 3 H2 Oð1Þ 4 2 ! FeðOHÞ3ðsÞ þ 2H2 SO4ðaqÞ

The actual oxidation process is considerably more complicated. Pyrite and related sulfide minerals contain both Fe and S in reduced oxidation states. When exposed to oxygen and water the sulfur moiety is oxidized first, releasing Fe2þ and sulfuric acid to solution [Reaction (2)]. The rate of oxidation is dependent on environmental factors like temperature, pH, Eh, and relative humidity as well as mineral surface area and microbial catalysis. 1 FeS2ðsÞ þ 3 O2ðgÞ þ H2 Oð1Þ ! FeðaqÞ 2þ 2 þ 2SO4ðaqÞ 2 þ 2HðaqÞ þ

Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001582 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

ð2Þ

Reaction (2) is most important in the initial stages of mine drainage generation and can be either strictly abiotic or mediated by contact with sulfur-oxidizing bacteria.[4] The Fe2þ released by pyrite decomposition is rapidly oxidized by oxygen at pH >3 as per Reaction (3).

What Causes Acid Mine Drainage? Mine drainage results from the oxidation of sulfide minerals such as pyrite (cubic FeS2), marcasite (orthorhombic FeS2), pyrrhotite (Fe1XS), chalcopyrite (CuFeS2) and arsenopyrite (FeAsS). These minerals are commonly found in coal and ore deposits and are stable until exposed to oxygen and water. Their oxidation causes the release of metals and the production of sulfuric acid. This process can occur as a form of natural mineral weathering but is exacerbated by mining because of the sudden, large-scale exposure of unweathered rock to atmospheric conditions.

ð1Þ

FeðaqÞ 2þ þ

1 1 O2ðgÞ þ HðaqÞ þ ! FeðaqÞ 3þ þ H2 OðIÞ 4 2 ð3Þ

If acidity generated by Reaction (2) exceeds the buffering capacity of the system, the pH eventually decreases. Below pH 3, Fe3þ solubility increases and a second mechanism of pyrite oxidation becomes important[5] [Reaction (4)]. FeS2ðsÞ þ 14FeðaqÞ 3þ þ 8H2 OðIÞ ! 15FeðaqÞ 2þ þ 2SO4ðaqÞ 2 þ 16HðaqÞ þ

ð4Þ 1

2

Acid Mine Drainage

Table 1 Summary of mine drainage from 101 bituminous coal mine sites in Pennsylvania Range pH

Median

Mean

2.7–7.3

5.2

3.6

Fe (mg=L)

0.16–512.0

43.0

58.9

Al (mg=L)

0.01–108.0

1.3

9.8

Mn (mg=L)

0.12–74.0

2.2

6.2

SO4 (mg=L)

120–2000

580.0

711.2

From Cravotta, C., III. USGS: Lemoyne, PA, 2001.

In this case, pyrite is oxidized by Fe3þ resulting in the generation of even greater acidity than when oxygen is the primary oxidant. Pyrite decomposition is thus controlled by the rate at which Fe2þ is converted to Fe3þ at low pH.[6] At pH 6.5) environments. Schwertmannite is commonly found in drainage waters with pH ranging from 2.8 to 4.5 and with

Copyright © 2006 by Taylor & Francis

Fig. 1 Mixing of acid mine drainage (at right) with a natural stream resulting in the formation of voluminous precipitates of iron minerals.

moderate to high sulfate contents. It may be the dominant phase controlling major and minor element activities in most acid mine drainage. Jarosite forms in more extreme environments with pH 100% 15-bar water content. Vitrisols with 20% C) have bulk density of about 0.2 g cm3. Soils with low C% (vitric horizons) have low clay content and often relatively high bulk density (>0.8 g cm3), but the soils have, in general, low bulk density characteristic of Andisols (predominately 2% Alo þ (1=2)Feo. The presence of vitric materials (various types of tephra) is also recognized by lowering this limit to as low as 0.4%, depending on the abundance of vitric materials. Most of the Icelandic Vitrisols meet the >0.4% Alo þ (1=2)Feo criterion for Andisols, but this limit seems arbitrary in Iceland and it can be doubted that it does provide meaningful separation of soils with abundance of vitric materials. On a world basis, vitric soils are poorly accounted for and recognized as weakly developed soils or parent material (tephra), with [Alo þ (1=2)Feo] < 0.4%. The uniqueness of these parent materials is recognized in some other classification systems such as for New Zealand with ‘‘Pumice Soils.’’[9]

ANDIC SOILS AND CRYOTURBATION Cryoturbation occurs with great intensity in Iceland.[1,10] The evidence is seen by cryoturbation in the soils and as various surface features such as ‘‘thufur’’ (hummocks) and a variety of active solifluction features. Cryoturbation in Iceland is enhanced by several factors. Water is pumped from shallow water table in wetlands. Freezethaw cycles are numerous and the soil temperature stays near 0 C for extensive periods which enhances stationary freezing front. The physical properties of Andosols are

108

important as well. Both saturated and unsaturated hydraulic conductivity of the soils are rapid which enhances water transfer to the freezing front. The great water holding capacity (often >40% at 15-bar suction in freely drained soils) also contributes water that freezes in the soil. An additional factor is the thixotropic nature of many soils, as this allows easy deformation of the soil by the ice formation.

SOIL EROSION The climate of Iceland is cold and subjected to periodic cold spells that exert stress on Icelandic ecosystems, as do ashfall events during volcanic eruptions. The vegetation of Iceland evolved in the absence of grazing animals. These factors render Icelandic ecosystems particularly sensitive to disturbance and reduced resilience caused by grazing and other land use. Soil erosion in Iceland has remained intense for the past 1200 years since the settlement by Vikings. A unique feature of soil erosion in Iceland is that the full depth of the soil mantle that has formed in eolian and tephra sediments is truncated, leaving barren surfaces (deserts) behind. A survey of erosion in Iceland has recently been published[11] showing the great extent and severity of soil erosion in Iceland. The Icelandic soils are extremely vulnerable to erosion, both by wind and water for many reasons. The formation of stable silt size clusters enhances saltation movement of soil particles making the soils susceptible to wind erosion. The lack of cohesion while wet and the thixotropic nature of the soils intensify water erosion and landslides. These andic properties that have reduced resistance against erosion are important contributing factors to the severe soil erosion in Iceland.

CONCLUSIONS A classification scheme specific to Iceland has been developed, and it has been used to produce a new 1:500,000 soil map.[1] Icelandic soils are primarily Andosols and Vitrisols according to the Icelandic scheme, which are both classified as Andisols according to Soil Taxonomy. The Icelandic soil environment is characterized by cold humid climate and a steady flux of vitric materials to the surface which, together


Andisols in Iceland

with drainage, are major factors dominating soil formation in Iceland. The extensive Icelandic desert environments are unique considering the humid climate of Iceland. The soils of the deserts are characterized by basaltic vitric materials and are termed Vitrisols. The andic soil characteristics of the Andisols, such as high water holding capacity, rapid hydraulic conductivity, thixotropy, and stable silt-sized aggregates, have important implications for intensive cryoturbation and soil erosion in Iceland.

REFERENCES 1. Arnalds, O. Volcanic soils of Iceland. In Volcanic Soil Resources. Occurrence, Development, and Properties; Arnalds, O., Stahr, K., Eds.; Catena Special Issue, Elsevier: Amsterdam, in press. 2. Quantin, P. Volcanic soils of France. In Volcanic Soil Resources. Occurrence, Development, and Properties; Arnalds, O., Stahr, K., Eds.; Catena Special Issue, Elsevier: Amsterdam, in press. 3. Arnalds, O. Sandy deserts of Iceland. J. Arid Environ. 2001, 47 (3), 359–371. 4. FAO. World Reference Base for Soil Resources. World Soil Resources Reports; FAO: Rome, 1998; Vol. 84. 5. Soil Survey Staff. Keys to Soil Taxonomy, 8th Ed.; USDA-NRCS: Washington, DC, 1998. 6. Parfitt, R.L.; Kimble, J.M. Conditions for formation of allophane in soils. Soil Sci. Soc. Am. J. 1989, 53 (3), 971–977. 7. Arnalds, O.; Hallmark, C.T.; Wilding, L.P. Andisols from four different regions of Iceland. Soil Sci. Soc. Am. J. 1995, 58 (1), 161–169. 8. Arnalds, O.; Kimble, J. Andisols of deserts in Iceland. Soil Sci. Soc. Am. J. 2001, 65 (6), 1778–1786. 9. Hewitt, A.E. New Zealand Soil Classification, 2nd Ed.; Landcare Research Science Series; Manaaki Whenau Press: Lincoln, New Zealand, 1998; Vol. 1. 10. Van Vliet-Lanoe, B.; Bourgeois, O.; Dauteuil, O. Thufur formation in northern Iceland and its relation to Holocene climate change. Permafr. Periglac. Process. 1998, 9 (4), 347–365. 11. Arnalds, O.; Thorarinsdottir, E.F.; Metusalemsson, S.; Jonsson, A.; Gretarsson, E.; Arnason, A. Soil Erosion in Iceland; Soil Conservation Service and Agricultural Research Institute: Reykjavik, Iceland, 2001. Translated from book published in Icelandic in 1997. Available on www.rala.is=desert.

Animals and Ecosystem Functioning Alan J. Franzluebbers United States Department of Agriculture-Agricultural Research Service (USDA-ARS), Watkinsville, Georgia, U.S.A.

INTRODUCTION Soil animals (i.e., fauna) are represented by a diverse array of creatures living in or on soil for at least a part of their life cycle. Many animals have influences on soil properties, but should not be considered soil dwellers since only a minor portion of their life cycle is spent in the soil (Fig. 1). Based on body size, soil animals can be divided into three categories: 1. microfauna (2 mm width) including millipedes, spiders, ants, beetles, and earthworms Soil animals can also be classified according to where they inhabit the soil. The aquatic fauna (e.g., protozoa, rotifers, tartigrades, and some nematodes) live primarily in the water-filled pore spaces and surface water films covering soil particles. Earthworms are divided into species that occupy the surface litter of soil (epigeic), that are found in the upper soil layers (endogeic), or that burrow deep into the soil profile (anecic). A further classification of five groups of soil animals is based on feeding activity, which can be useful in distinguishing how different groups affect soil ecosystem functions: 1. Carnivores feed on other animals. This group can be subdivided into: i) predators (e.g., centipedes, spiders, ground beetles, scorpions, ants, and some nematodes), who normally engulf and digest their smaller prey and ii) parasites (e.g., some flies, wasps, and nematodes), who feed on or within their typically larger host organism. 2. Phytophages feed on living plant materials, including those that feed on above-ground vegetation (e.g., snails and butterfly larvae), roots (e.g., some nematodes, fly larvae, beetle larvae, rootworms, and cicadas), and woody materials (e.g., some termites and beetle larvae). Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001731 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

3. Saprophages feed on dead and decaying organic material and include many of the earthworms, enchytraeids, millipedes, dung beetles, and collembola (or springtails). Saprophages are often referred to as scavengers, debris-feeders, or detritivores. 4. Microphytic feeders consume bacteria, fungi, algae, and lichens. Typical microphytic feeders include mites, collembola, ants, termites, nematodes, and protozoa. 5. Miscellaneous feeders are not restrictive in their diet and consume a range of the previously mentioned sources of food. This group includes certain species of nematodes, mites, collembola, and fly larvae. The arrangement of these feeding groups can be visualized as a soil food web with multiple trophic levels, beginning with the autotrophic flora (Fig. 2). Trophic levels describe the order in the food chain. The first trophic level is composed of photosynthetic organisms, including plants, algae, and cyanobacteria, which fix CO2 from the atmosphere into organic compounds. Organisms that consume the photosynthesizers are in the second trophic level, which includes bacteria, actinomycetes, fungi, root-feeding nematodes and insects, and plant pathogens and parasites. The third trophic level feeds on the second trophic level, including many of the dominant soil animals, including bacterial- and fungal-feeding arthropods, nematodes, and protozoa. The soil food web can be continued to include various vertebrates, including amphibians, reptiles, and mammals.

SPATIAL DISTRIBUTION OF SOIL ANIMALS Soil animals are not uniformly distributed in soil. Unlike the soil microflora, which could be considered ubiquitous, the proliferation of soil animal communities is more sensitive to environmental disturbances and ecological interactions. Gross climatic differences afford opportunities for unique assemblages of organisms. Even within a specific climatic region, large differences occur in the community of organisms present based upon type of vegetation, soil, availability 109

110

Animals and Ecosystem Functioning

Fig. 1 Categories of soil animals defined according to degree of presence in soil, as illustrated by some insect groups. (From Ref.[1].)

of water, land use, and presence of xenobiotics. Within the confines of a seemingly uniform pedon, ‘‘hot spots’’ of soil organism activity can be isolated based on localized availability of resources and environmental conditions (Fig. 3).

2.

3. INFLUENCE OF SOIL ANIMALS ON SOIL FUNCTIONS Decomposition and Nutrient Cycling Soil animals work directly and indirectly with the soil microflora (i.e., bacteria, actinomycetes, fungi, and algae) to decompose organic matter and mineralize nutrients.[3] The primary consumers of organic materials are the soil microflora. Soil animals, like many of the microflora, are heterotrophs and therefore consume organic materials to gain energy for growth and activity. Soil animals make important contributions to decomposition by

4. 5.

6.

7.

activities of other organisms, especially microorganisms; consuming resistant plant materials that would decompose slowly otherwise, such as wood, roots, and dung, and transforming these materials into more decomposable constituents; dispersing soil microorganisms (i.e., inoculation) within the soil profile by transporting them on their bodies and through their intestinal tracts; creating burrows in soil to increase aeration, which stimulates microbial activity; transporting organic materials from the soil surface to deeper in the soil profile, thereby improving environmental conditions for decomposition and increasing biological interactions deeper in the soil profile; consuming bacteria and fungi, thereby releasing nutrients and stimulating the regeneration of microbial populations; and providing unique food sources themselves for consumption by other soil fauna and microflora.

1. shredding organic materials, thereby exposing a greater surface area for enhancing the Water Cycling

Fig. 2 Generalized diagram of a soil food web.


Soil animals are active participants in the formation of soil structure, which is an important characteristic that influences water infiltration, soil water retention, and percolation.[4] The biochemical activity of soil organisms transforms organic materials into soil-stabilizing cementing agents, which bind the primary soil particles (i.e., sand, silt, and clay) into aggregates. In addition, the burrowing activity of soil animals creates larger pores alongside water-stable aggregates to increase total porosity of soil, which aids water flow without decreasing overall water retention capacity and improves the plant rooting environment. Both aggregates and porosity are important components of soil structure. Poor soil structure due to

Animals and Ecosystem Functioning

111

Fig. 3 Key locations of soil organism activity. (From Ref.[2].)

disruption of aggregates, which fills pores with disaggregated primary particles and causes crusting of the soil surface, results in more rainfall that runs off land (i.e., less infiltration), potentially carrying with it sediment, nutrients, and pesticides that can contaminate surface waters. Reduced infiltration with poor aggregation reduces available water for plant growth (i.e., reduces net primary productivity and the potential to fix atmospheric CO2) and reduces percolation of water through the soil profile, essential for purification and recharge of groundwater. Those animals that create burrows in soil also create conduits for water movement through the soil profile. These biopores can be important for improving water percolation and improving rooting below claypans and other restrictive soil layers. Many different soil animals deposit fecal pellets, which become stable soil aggregates when the organic material is mixed with soil mineral particles. These aggregates are able to retain more water because of the high water-holding capacity of soil organic matter.

Pest Control Intense competition among soil organisms keeps an ecosystem healthy by preventing one organism from becoming dominant. Potential plant pathogens,


such as root-feeding nematodes, are often held below damaging levels because of consumption by predatory nematodes and arthropods. With a healthy food web rich in species diversity, the predatory activity of many arthropods can keep crop pests below economic thresholds.

Impact of Key Soil Animals Earthworms Earthworms are well-known soil animals inhabiting many environments, most prominently found in moist-temperate ecosystems. As earthworms ingest organic materials and mineral particles, they excrete waste as casts, which are a particular type of soil aggregate that is rich with organic matter and mineralizable nutrients. It is estimated that a healthy population of earthworms can consume and aerate a 15 cm surface of soil within one or two decades. Anecic or deepburrowing species of earthworms can create relatively permanent vertical channels for improving root growth and water transport. Important attributes of earthworm activities are increased surface soil porosity, enhanced water infiltration and nutrient cycling, and distribution of organic matter within the soil profile to increase soil microbial activity.

112

Termites Termites are important soil animals in grasslands and forests of tropical and subtropical regions. They often build mounds by excavating subsoil and depositing it above ground to build a city of activity with a complex social system. Termites are able to decompose cellulose in wood because they harbor various microorganisms (protozoa, bacteria, or fungi) to aid in decomposition. Better drainage and aeration of termite mounds may be beneficial to nearby plant growth in soils with a high water table. Stable macrochannels created by termites can improve water infiltration into soils that otherwise would form impermeable surface crusts. Protozoa Protozoa are single-celled animals that generally consume bacteria and soluble organic matter. Protozoa are more numerous in marine and freshwater environments, but do occur widely in water films of many soils.[5] Their principal soil function is predation on soil bacteria, which releases nutrients for potential plant uptake; increases decomposition and soil aggregation by stimulating their bacterial prey; and prevents some bacterial pathogens from developing on plant roots.

SOIL BIODIVERSITY There has been a great deal discovered about soils and the organisms that live in them, yet it is estimated that 1 m in height and >2 m diameter.[2] Ant nests consist of underground, branched networks of galleries and chambers. Surficial chambers are connected to lower chambers by vertical galleries with branching lateral galleries. Galleries and chambers vary in size and number, depending upon the species of ant. For example, Lasius neoniger, an abundant ant species in temperate North America, constructs tubular galleries of 1.5–5.0 mm diameter and chambers of 10– 20 mm diameter and 30–50 mm length. The volume of L. neoniger nests ranges from 20–250 cm3 and the nests are confined to the upper 70 cm of soil.[3] Other species construct nests to depths ranging from 50 cm to greater than several meters, depending upon species-specific behavior, soil type, and landscape position. Soil profile mixing, texture, physical and chemical property modification of mound soils, soil macroporosity, and geomorphological attributes of ant nest mounds vary with species-specific colony longevity, body size, and numbers of workers of a colony, soil type, and landscape position. The pedturbation effects of ants therefore depend upon the species composition of the ant community, geomorphic history, soil properties, and topographic position of a landscape unit. Because most studies concerning the effects of ants on soils have focused on one or two species, a comprehensive analysis of the combined effects of all ant species on the soils of an ecosystem cannot be made.

In areas that are periodically flooded or where the water table is close to the surface, some species of soil-nesting ants build mounds that create favorable microhabitats for themselves and also a habitat for some species of plants that are confined to the aerated soils of the ant mounds. Soil-nesting ants create hummock microtopography in some wet meadow fens and tropical wet savannas.[4] In the Chaco region of South America (parts of Paraguay, Bolivia, Argentina, and Brazil), nest mounds of Camponotus punctulatus occur at densities of between 200 and 1000 mounds=ha. These conical mounds average a height of 0.62 cm (with a maximum of 1.85 m ) with a mean basal diameter of 1.2 m. The mound soils are lighter textured than surrounding soils, reflecting the amount of materials transported from surrounding subsurface soil during mound construction.[5] Formica podzolica mounds in a Montana fen are thought to contribute to the hummock-hollow microtopography of peat lands. Abandoned F. podzolica mounds provide drier, warmer microsites that are enriched with some soil nutrients.[4] The mounds of L. flavus contribute to the microtopography of some European grasslands and salt marshes.[6] Mima-type earth mounds up to a height of 1.5 m with a diameter of 20 m in Buenos Aires Province, Argentina, are produced by horizontal translocation of soil to the colony sites of black fire ants, Solenopsis richteri. Continued occupation of the mounds by successive generations of ants gradually increases the size of the mounds to mima-type size.[7] Ants (Formica spp. and Myrmica spp.) are important agents in the process of development and maintenance of hummock microtopography of subarctic peatlands. Hummock retrogression is accelerated by the tunneling activity of ants.[8]


HETEROGENEITY OF PHYSICAL AND CHEMICAL PROPERTIES OF THE SOIL Many species of ants alter the texture and chemistry of the soil in the nest mounds. The nutrients most 117

118

frequently reported to be at higher concentrations in ant mound soils include nitrogen, phosphorus, potassium, calcium, magnesium, manganese, and iron.[9] The effect of soil-nesting ants on soil nutrient patchiness and on vegetation varies as a function of landscape position, soil type, and the biology of the ant species. Nutrient enrichment of mound soils has been reported for several species of seed-harvesting ants and omnivorous species of ants that collect seeds, prey on insects, or collect insect carrion. Species of soilnesting ants that enrich the nutrient content of mound soils are characterized by relatively long-lived colonies (>5 years) and the behavior of depositing chaff and unwanted insect parts on and around the nest mound or disk. Nutrient enrichment of mound soils by a species may not occur on all soils on a watershed or landscape. For example, Pogonomyrmex rugosus nest disks in desert shrubland and mixed shrub-grassland were nutrient enriched, but the nest disks of this species in a piedmont grassland were not nutrient enriched.[10] Formica spp. mounds in forest were nutrient enriched, but Formica spp. mounds in meadows and grasslands were not.[11] The variability in soil nutrient enrichment of ant mounds has been documented in several species of leaf-cutting ants. In remnant Cerrado (woodlandsavanna), Brazil, leaf-cutting ants (Atta spp.) had no detectable effect on nutrient enrichment.[12] In northern Patagonia, soils associated with the leaf-cutting ant, Acromyrmex lobicornis, had higher concentrations of nitrogen, phosphorus, and organic matter than reference soils.[13] The location of nutrient-rich organic refuse produced by leaf-cutting ant colonies varies among species. A. cephalotes deposit organic refuse in subterranean chambers, whereas A. colombica place organic refuse on the soil surface near the nest. The location of organic refuse is a major factor affecting nutrient concentrations and the composition, abundance, and activity of soil microflora and microfauna.[14] In the Orinoco Llanos savanna, Venezuela, A. laevigata nests had higher concentrations of nitrogen, magnesium, calcium, and organic carbon, but other soil nutrients and properties were not affected by ant mounds.[15] In an Australia vertisol, ant nest soils had greater concentration of coarse and particulate organic matter, lower fine particulate soil organic matter (SOM) =coarse particulate SOM ratios, larger sand content, and lower clay content than surrounding soils.[16] Nutrient enrichment of nest mound soils of funnel ants (A. barbigula) was attributed to entrapment of organic materials around the nest entrances. Re-excavation of nest chambers after rainfall buries trapped litter, resulting in higher concentrations of nitrogen, organic matter, and some cations compared to nest-free soils.[17] In humid tropical savanna, ant mounds of


Ants

Camponotus spp. had higher clay and coarse sand content than surrounding soils.[9] Even exotic or alien species of ants change the chemical and physical properties of nest mound soils. Mounds of imported fire ants (S. invicta) had higher concentrations of clay, phosphorus, and potassium, and lower concentration of soil organic matter than reference soils. The effect of S. invicta on calcium concentrations relative to reference soils was dependent upon the characteristics of the unmodified soil.[18] Ants change the nutrient concentration of mound soils, but the physical and chemical properties of mound soils can also affect mineralization processes. Nitrogen mineralization rates were reduced in nest mound soils in moss-sedge, sedge, and alder peat habitats.[19]

SOIL TURNOVER The longevity and turnover rates of nests and nest mounds of ants in a community frequently follow a distribution gradient from high turnover (10 years). The importance of ants in the transport of subsurface horizon materials to the surface varies with the density and diversity of the ant community on a landscape unit. In Chihuahuan Desert grasslands, soil-nesting ants are an order of magnitude more abundant on sand and sandy loam soils than on fine-textured soils. Ants were estimated to move between 21.3 and 85.8 kg=ha=yr on sandy and sandy loam soils and between 0.1 and 3.4 kg=ha=yr on clay and clay-loam soils.[20] The estimated annual soil turnover by ants in an Atriplex vesicaria shrubland in the semi-arid region of Australia was 350– 420 kg=ha=yr.[9] Soil that is excavated by ants in the construction of galleries and chambers and deposited on mounds around nest entrances is generally eroded by water and wind within a year unless the mound is protected from raindrop splash erosion by gravel, stones, or wood fragments. Nest mound soils may be replenished by the belowground expansion of galleries and chambers. Ant nest mounds in sparsely vegetated arid regions are prone to wind erosion. On an Australian aeolian soil, funnel ants’ (Aphaenogaster barbigula) nests were active for approximately nine months and the ants changed location approximately twice per year. Soil transport was estimated to be 33.6 kg=ha, and it was estimated that 92% of the soil volume would be turned over by these ants in 100 years.[21] In Western Australia, ant communities on gray soils of semi-arid woodlands were estimated to turnover 46.5 kg ha=yr and on yellow soils, the soilnesting ant community was estimated to turnover 22.3 kg=ha=yr.[22] In a humid savanna environment, one abundant ant species, Paltothyreus tarsatus, was

Ants

estimated to transport approximately 30 g=m2=yr of sand particles and soil aggregates. This ant species increased the concentrations of clay, carbon, iron oxides, and coarse sand in the A horizon.[9] The amount of soil transported to the surface by Pognomyrmex occidentalis in pinon-juniper woodland and ponderosa pine forest was estimated to be 650 kg=ha.[23] Soil turnover by the ant community in New England forest soil was estimated to be over 50 kg=ha=yr. It was concluded that the translocation of B-horizon materials to the soil surface by soilnesting ants was an important process in podzol formation in New England forest soils.[24] Some long-lived species of soil-nesting ants relocate their nests one or more times a year. Construction of new nests results in the transport of a volume of soil equal to the volume of galleries and chambers to the soil surface. Most of that soil originates in lower soil horizons and contributes to soil profile homogenization. The relocation of nests by some species of ants results in lower estimates of soil turnover than occurs in some environments.

SOIL WATER RELATIONS The structure of nests of soil-nesting ants provides extensive macroporosity to the soil in which the nests are constructed. The macropores constructed by ants affect rates of infiltration and rates of percolation. In some environments, extremely high densities of nest entrances can have a dramatic effect on infiltration. In semi-arid Western Australia, ant biopores were found to transmit water down the soil profile only when the soil was saturated and water was ponding on the surface.[25] On aeolian sand soils in Australian semi-arid woodland, densities of nest entrances of funnel ants (A. barbigula) were estimated at 88,000 per hectare. Steady-state water infiltration on soils with nest entrances averaged 23.3 mm=min, in comparison to an infiltration rate of 5.9 mm=min on nestentrance-free soil.[26] In semi-arid woodland of Eastern Australia on red earth soil, ponded steady-state infiltration averaged 1026 mm=hr on soil with nest entrances of A. barbigula, but only 120 mm=hr on soils without nest entrances.[27] Bulk flow along nest galleries provides an important route of recharge of deep soil moisture in arid and semi-arid environments. Ant gallery macropores are not always avenues for bulk flow. In a study of a mesic Typic Quartzipsamment, there was no preferential flow down ant galleries. The lack of an effect on hydraulic conductivity was attributed to the sandy soil.[28] In another study of a sandy soil, the estimated saturated soil matrix hydraulic conductivity of nest burrows was approximately


119

eight times smaller than that of the bulk sandy soil. This reduction in hydraulic conductivity was attributed to the ants’ in-filling of gallery walls with fine materials.[28]

EFFECTS ON OTHER SOIL BIOTA Soil around relatively long-lived ant colonies may be enriched with microflora, microfauna, and mesofauna. The soils of nest disks of western harvester ants, P. occidentalis, are enriched with vesicular-arbuscular mycorrhizal fungi.[29] In areas of North America dominated by the red imported fire ant, S. invicta, the species composition and abundance of soil yeast within mounds are altered by changes in soil properties produced by fire ants.[30] Mound soils of F. aquilonia are dominated by bacteria-feeding microfauna and have a higher microbial biomass than the surrounding soils.[31] Species-specific differences in the effect of ants on soil microflora of mounds are related to the feeding strategies of the species and nest architecture. Three ant species, M. scabrinodis, L. niger, and L. flavus, differ greatly in foraging strategies and methods of mound construction. Microbial functional diversity and evenness were higher in mound soils of M. scabrinodis and L. niger than in reference soils but were not different from reference soils in the mounds of L. flavus. Different functional groups of microorganisms were activated in the mounds of the different species. Carbon mineralization was higher in mound soils of all three species.[32]

CONCLUSIONS Ants contribute to heterogeneity in soil properties by the construction of subterranean nests and by accumulating organic materials in and around nests. Construction and maintenance of nests affect soil turnover, macroporosity, and mixing of soil profiles. Foraging and food processing behaviors affect soil nutrient concentrations and soil microflora and microfauna. The magnitude of the effects of ants on soils is dependent upon soil type and topographic position.

REFERENCES 1. Holldobler, B.; Wilson, E.O. The Ants; The Belknap Press of Harvard University Press: Cambridge, 1990. 2. Green, W.P.; Pettry, D.E.; Switzer, R.E. Formicarious pedons, the initial effect of mound-building ants on soils. Soil Survey Horizons 1995, 39 (2), 33–44. 3. Wang, D.; McSweeney, K.; Lowery, B.; Norman, J.M. Nest structure of the ant Lasius neoniger emery and

120

4.

5.

6. 7.

8.

9.

10.

11.

12.

13.

14.

15.

16.

17.

18.

Ants

its implications to soil modification. Geoderma 1995, 66 (3–4), 259–272. Lesica, P.; Kannowski, P.B. Ants create hummocks and alter structure and vegetation of a montana fen. Am. Midl. Nat. 1998, 139 (1), 58–68. Pire, E.F.; Torres, P.F.; Romagnoli, O.D.; Lewis, J.P. The significance of ant-hills in depressed areas of the Great Chaco. Rev. Biol. Trop. 1991, 39 (1), 71–76. King, T.J. Ant-hills and grassland history. J. Biogeog. 1981, 8, 329–334. Cox, G.W.; Mills, J.N.; Ellis, B.A. Fire ants (Hymenoptera : Formicidae) as major agents of landscape development. Environ. Entomol. 1992, 21 (2), 281–286. Luken, J.O.; Billings, W.D. Hummock-dwelling ants and the cycling of microtopography in an Alaskan peatland. Can. Field Nat. 1986, 100 (1), 69–73. Lobry de Bruyn, L.A.; Conacher, A.J. The role of ants and termites in soil modification a review. Aust. J. Soil Res. 1990, 28 (1), 55–95. Whitford, W.G.; DiMarco, R. Variability in soils and vegetation associated with harvester ant (Pogonomyrmex rugosus) nests on a Chihuahuan desert watershed. Biol. Fertil. Soi. 1995, 20, 169–173. Culver, D.C.; Beattie, A.J. Effects of ant mounds on soil chemistry and vegetation patterns in a Colorado montane meadow. Ecology 1983, 64 (3), 485–492. Schoereder, J.H.; Howse, P.E. Do trees benefit from nutrient rich patches created by leaf-cutting ants? Stud. Neotrop. Fauna Environ. 1998, 33 (2–3), 111–115. FarjiBrener, A.G.; Ghermandi, L. Influence of nests of leaf-cutting ants on plant species diversity in road verges of Northern Patagonia. J. Veg. Sci. 2000, 11 (3), 453–460. FarjiBrener, A.G.; Medina, C.A. The importance of where to dump the refuse: seed banks and fine roots in nests of the leaf-cutting ants. Atta cephalotes A. colombica. Biotropica 2000, 32 (1), 120–126. Brener, A.G.F.; Silva, J.F. Leaf-cutting ants and forest groves in a tropical parkland savanna of Venezuela: facilitated succession. J. Trop. Ecol. 1995, 11 (4), 651– 669. Hulugalle, N.R. Effects of ant hills on soil physical properties of a vertisol. Pedobiology 1995, 39 (1), 34–41. Eldridge, D.J.; Myers, C.A. Enhancement of soil nutrients around nest entrances of the funnel ant Aphaenogaster barbigula (Myrmicinae) in semi-arid Eastern Australia. Aust. J. Soil Res. 1998, 36 (6), 1009–1017. Green, W.P.; Pettry, D.E.; Switzer, R.E. Impact of imported fire ants on the texture and fertility of Mississippi soils. Comm. Soil Sci. Plant Anal. 1998, 29 (3–4), 447–457.


19. Petal, J. The influence of ants on carbon and nitrogen mineralization in drained fen soils. Appl. Soil Ecol. 1998, 9 (1–3), 271–275. 20. Whitford, W.G.; Forbes, G.S.; Kerley, G.I. Diversity, spatial variability, and functional roles of invertebrates in desert grassland ecosystems. In The Desert Grassland; McClaran, M.P., Van Devender, T.R., Eds.; University of Arizona Press: Tucson, 1995; 152–195. 21. Eldridge, D.J.; Pickard, J. Effects of ants on sandy soils in semiarid Eastern Australia. 2. Relocation of nest entrances and consequences for bioturbation. Aust. J. Soil Res. 1994, 32 (2), 323–333. 22. Lobry de Bruyn, L.A.; Conacher, A.J. The bioturbation activity of ants in agricultural and naturally vegetated habitats in semiarid environments. Aust. J. Soil Res. 1994, 32 (3), 555–570. 23. Carlson, S.T.; Whitford, W.G. Ant mound influence on vegetation and soils in a semiarid mountain ecosystem. Am. Midl. Nat. 1991, 126 (1), 125–139. 24. Lyford, W.H. Importance of Ants to Brown Podzolic Soil Genesis in New England; Harvard Forest Paper No. 7; 1963; 1–18. 25. Lobry de Bruyn, L.A.; Conacher, A.J. The effect of ant biopores on water infiltration in soils in undisturbed brushland and farmland in a semi-arid environment. Peodobiology 1994, 38 (3), 193–207. 26. Eldridge, D.J. Effect of ants on sandy soils in semiarid Eastern Australia: local distribution of nest entrances and their effect on infiltration of water. Aust. J. Soil Res. 1993, 31 (4), 509–518. 27. Eldridge, D.J. Nests of ants and termites influence infiltration in a semiarid woodland. Pedobiology 1994, 38 (6), 481–492. 28. Wang, D.; Lowery, B.; Norman, J.M.; McSweeney, K. Ant burrow effects on water flow and soil hydraulic properties of sparta sand. Soil Till. Res. 1996, 37 (2–3), 83–93. 29. Friese, C.F.; Allen, M.F. The interaction of harvester ants and vesicular arbuscular mycorrhizal fungi in a patchy semiarid environment: the effects of mound structure on fungal dispersion and establishment. Funct. Ecol. 1993, 7 (1), 13–20. 30. Ba, A.S.; Phillips, S.A.; Anderson, J.T. Yeasts in mound soil of the red imported fire ant. Mycol. Res. 2000, 104 (8), 966–973. 31. Laasko, J.; Setala, H. Composition and trophic structure of detrital food web in ant nest mounds of Formica aquilonia and in the surrounding forest soil. Oikos 1998, 81 (2), 266–278. 32. Dauber, J.; Wolters, V. Microbial activity and functional diversity in mounds of three ant species. Soil Biol. Biochem. 2000, 32 (1), 93–99.

Arab Traditional Soil Classification: A Moroccan Case Mohamed Sabir Hassan Ben Jelloun National School of Forest Engineers, Tabriquet, Sale, Morocco

INTRODUCTION Soil science, a relatively new discipline, compared to other fields, such as botany, biology, zoology, and mineralogy, is lacking an internationally accepted taxonomic system. Traditionally, a local scientific system of soil classification was lacking in most countries of the Arab world. However, soils were mostly designated by vernacular names that stem from local agricultural practices and soil use.

BACKGROUND Different systems of soil classification have been developed throughout the world.[1] In the Maghreb Arab countries (Morocco, Algeria, and Tunisia), the first efforts related to soil classification were initiated in 1934 by Del Villar, who was the president of the Mediterranean subcommission of the 1st International Society of Soil Science.[2] His principal and most important pedological studies were conducted in Morocco. At present, vernacular names are still used among farmers and even among agricultural scientists, and detailed soil classification system are inspirations from the ones developed in western Europe, Canada, and the United States of America, especially the systems of soil classification adopted by the Commission of Pedology and Cartography of Soil of France (CPCS, 1967),[3] FAO=UNESCO. Soils of the World (ELSEVIER=ISRIC, 1989),[4] Soil Taxonomy (1975)[5] and the World Reference Base (1994).[6] The object of this entry is to present a brief review about the soil classification scheme used in some areas of Morocco.

ANCIENT ARAB SOIL CLASSIFICATION Knowledge about land and agricultural practices stems from before Roman times. Men in ancient times have already used many of our farming practices nowadays (manuring, liming, and crop rotations with legumes).[7] Most of the actual knowledge farmers used during the long period from the fall of Roman civilization Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120017330 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

to the French Revolution and for some time afterward was the local knowledge of farmers. At the time where the people of Europe were disorganized and lived in a dark age of diseases, famine, and war for more than a thousand years, the Arabian culture flourished in the Near East, northern Africa, and southern Spain, where farming was reasonably good, especially under irrigation. Some of soil classification systems and farming practices were well explained in the handbook of agriculture prepared by Ibn-al-Awam,[7] with a Moorish scholarship, in the 12th century. Ibn-al-Awam’s[8] book of agriculture classified different land types according to their agricultural qualities as follows:

Black earth: A warm earth that gives high agricultural yields;

Red earth: A moderately warm earth with moderate agricultural yields;

Yellow earth and white earth: A cold earth that gives low agricultural yields;

Dry earth: It has two species or divisions: – Sandy earth: It has a low fertility level; – Muddy earth: It has a marly–clay texture, it is sticky and plastic, and it becomes very hard and very compact when dry.

MOROCCAN VERNACULAR SOIL NAMES Soil classification was introduced for the first time to the Mediterranean region in 1925 through the efforts of Spanish pedologist Del Villar (Table 1).[9] Previously, most of the soil types were designated by their color, texture, structure, degree of fertility, and water-holding capacity despite the fact that their meaning may differ from a region to another. In the northern humid regions, soils are distinguished by their fertility. Contrariwise, in the southern arid regions, soils are distinguished by their water-holding capacity following irrigation. Among the Maghreb West Arab countries, Morocco, Algeria, and Tunisia show climatic, geologic, and geomorphologic similarities.[10] Therefore, some of the vernacular soil names (Table 2) used in Morocco may, to some extent, be encountered in Algeria and Tunisi. 121

122

Arab Traditional Soil Classification: A Moroccan Case

Table 1 Del Villar soil classification system Soil types 1-Homocyclic type

Soil cycles 1.1. Sialferric cycle: Soil dynamics are dominated by colloidal silicon and aluminum and by ferrous sesquioxides.

Soil sectors S1. Oxyhumic sector: Unsaturated soil with acid humus. Leaching or epigenic metabolism is strong (e.g., gley soil). S2. Siallic sector: Humus is not acidic, the ratio SiO2=Al2O3 > 2, pH 2; Fe2O3 is high. S4. Allitic sector: SiO2=Al2O3 is very low. Leaching is important. Podzolization is possible. Ferruginous accumulation is possible (laterite soil).

1.2. Calcareous cycle: Pedogenic processes related to calcareous parent rock or carbonate-rich material are dominant. Leaching of carbonate may give K horizon at depth. Distinguished soils are: Soil with AC profile; Soil with AKC profile (rendzina soil, terra-rossa soil). 1.3. Sodic cycle: Pedogenic processes are dominated by sodium (Na). S1. Saline sector: Na is in chloride form to which other soluble salts and gypsum are added. Saline soils may be epigenic or hypogenic. Epigenic metabolism of Na (leaching) is dominant (e.g., coast marshes). Hypogenic metabolism: Salts are lifted from below to the surface. Examples are: Thermal sources; Salty lagoons; Black saline soil: gley solontchak soil; Local deposited salt sebkhas of Arab countries; Saline soil with surface crust; Soil with salt at depth (infrasaline soil); Soil with low salt concentration at surface (subsolontchak soil). (Continued)



123

Table 1 Del Villar soil classification system (Continued) Soil types

Soil cycles

Soil sectors S2. Alkaline sector: Hypogenic metabolism process is low, but endogenic process resulting from precipitation, epigenic process, and percolation is dominant (solonetz soil).

1-Heterocyclic type: Pedogenic metabolism responsible for soil type is not based on known chemical process. Chemical characteristics are mixed or variable; it is the regime which is responsible of soil type individualization. Distinguished soils are: Nonsodic–hydro-epigenic soil (alluvial soil); Soil with subamphigenic or subbalanced regime (prairie soil); Soil with amphigenic or balanced regime (chernosium soil); Soil with hydro-hypogenic regime (clayey–gley soil); Soil with hydro-hypogenic calcification; Oligogenic soil (Thin A horizon); Ambrogenic or holohypogenic soil ¼ crypto-hypogenic and pheno-hypogenic soil (soils of desert climate. Below ground water table and lithic material are the dominant genetic factors.); Mixed transitional soil types.

In Morocco (Fig. 1), several local soil names are used. Region I: North Western Rif Mountains Ferich: A thin soil formed on pelitic rock material with high percentage of flysch plates. Rmel and ferich: Complex soil, mixture of sandy (rmel) and ferich materials. Abiad: White soil color (abiad means white), developed on marl and marly–flysch material. Hajar: Nonsoil with sandstone outcrops (hajar means rock outcrop and stones). Rmel or rmel jayef: Sandy soil (rmel means sand) with red-yellowish color. This soil has low fertility, sometimes has small rock grains, and is not hard when dry. Teine: Vertic soil with high expanding clay content and high water retention (teine means clay). Hamri: Red soil with high content of ferric oxides. Amlil: Lithic calcareous soil with generally white color. Sahl: Less developed soil on alluvial sediments. Ard kbira: Black soil with high fertility level and high agricultural productivity and is suitable for


all agricultural uses with low rock content and low water infiltration rate (ard kbira means big soil). Toiress: Brown soil with low fertility but still good crop production. Rock content of this soil is moderate. Fairly hard when dry with particular texture. Region II: Temara (South of Rabat on the Coast Side) Rmel: Sandy soil with brown color. Highly permeable, not hard when dry, has low percent of stones, and may reach 2-m depth. Hamri: Red soil with high content of ferric oxides in mixture with clay; it is weakly permeable, hard, has low percentage of stones, and may reach 1-m depth. El hassa: Grey soil with high percentage of stones, fairly permeable, and stones may reach 0.5-m depth. Biad: White-colored soil with low agricultural productivity and good water permeability. Hrach: Yellow-colored soil with low agricultural productivity, good water permeability, and some stones (hrach means rough).

124


Table 2 Vernacular soil names and their equivalent in soil taxonomy Vernacular soil names a

Soil taxonomy equivalent names a

Ferich (R1 ), Hajar (R1), Sgguine (R5 ), Salsale (R5), Azegzou (R7a)

Entisols

Rmel (R3a and R5)

Psaments (Entisols)

Rmel (R1), Abiad (R1), Amlil (R1), Sahl (R1)

Udepts (Inceptisols)

Rmel jayef (R1)

Mixed Entisol–Udepts

Biad (R2a), El Hrach (R5)

Xerocrepts (Inceptisols)

Hrach (R2 and R3)

Stony xerocrepts

Biadi (R3)

Xerocrepts over marl

Teine (R1)

Pelluderts (Vertisols) a

Tirs (R3), Tirst (R4 ), Tirste (R5)

Chromoxererts (Vertisols)

Hrach (R6a), Itikki or Hamri (R6), Tamakalte (R6), Terrist (R7)

Orthids (Aridisols)

Amarigh (R7)

Salorthids (Aridisols)

Azougakh (R7), Amerdoum mouharmel (R7)

Argids (Aridisols)

Ard Kbira (R1), Toires (R1)

Vertic Udolls (Mollisols)

Rtab (R4), Hamri (R5)

Xerolls (Mollisols)

Safra (R4)

Aquolls (Mollisols)

Rmel (R2 and R4), Hamri (R2, R3), and R5), Ahamri (R4)

Xeralfs (Alfisols)

a

R1, R2, R3, R4, R5, R6, and R7 are Moroccan regions where local soil names are used.

Region III: Romani (South East of Rabat) Tirs: Black to dark brown soil with high fertility, needs large quantities of water for irrigation, has good agricultural potential, contains high percentage of expanding clay and low stones, has good permeability, and is resistant to erosion. Hrach: Grayish soil that requires less water with medium agricultural potential, is slightly permeable, and has large particles and stones. Hamri: Red soil that has particles of loam and clay diameter in mixture with stones, needs much water for irrigation with very good agricultural potential. Biadi: White soil with dust and stone on the surface. Rmel: Brownish soil with sandy texture, high permeability, and low water-holding capacity, no stones, and deep. Region IV: Midelt (High Moulouya Region) Ahamri: Red soil that has clayey texture, compact with no stones, a high water-holding capacity, gets rapidly dry and warm, and has good crop production capacity. Tirst: This soil is black, has clayey texture, is less or more compact with no stones. Water-holding capacity of this soil is fairly good, and it is good soil for agriculture. It shows more resistance to dryness than hamri. On slopes, when dry, may


show wind erosion hazard. Parent rocks are sediments of 10-cm depth. Agricultural yields are of 1000 to 1200 kg=ha=yr (less than world average). Rtab: This is dusty soil with medium texture and no stones. It is 1 m thick, holds water for 20 to 30 days after showers or irrigation. Hrach: Hrach is dark brown compact soil with large particles and stones. It is 0.5 to 1 m thick, holds water for 8 to 10 days after showers or irrigation. Rmel: It has silver-gray color, light-sandy texture, and low water-holding capacity. Safra: Yellow calcareous soil (safra means yellow color) with clayey texture, hard, shows some water stagnation, and shallow (about 15 cm thick). Region V: Moulay Bouazza Tirste: Black soil with slightly stony–fine texture, low permeability, high water-holding capacity, shows little cracks when dry and good biomass production vegetation development. Hamri: It is a red compact soil, slightly stony texture, low to moderate water permeability, and cracks when dry. This soil needs much water for irrigation and shows medium potential for agriculture. El Hrach: This soil has reddish-gray color with stones in the surface layer, is fairly permeable, and requires low quantities of water for irrigation.


125

Fig. 1 Map of Morocco, geographical situation of different regions where traditional soil names were collected (RI: North Western Rif; RII: Temara; RIII: Romani; RIV: Midelt; RV: Moulay Bou Azza; RVI: Bou Malne Daddes; RVII: Erfound).

Rmel: Yellow soil of sandy texture, with no stones, and highly permeable. Sgguine: White-colored soil with hard stony–fine texture, permeability is low and agricultural potential is poor. Salsale: Gray soil with low stones, has friable consistence, low permeability, and poor quality. Region VI: Bou Malne de Daddes (arid and Saharan region) Rmel: It has yellow soil color, fine texture with no stones, not hard when dry, fertility is medium, and water-holding capacity is very low.


Hrach: Faint red soil, very hard stony–loamy texture, fertility is moderate, water-holding capacity is low, workability is difficult, sometimes has water stagnation, and agricultural potential is medium. Itikki or hamri: Red soil with very hard compact clayey texture, has a high water-holding capacity, its dryness is very slow, and shows a good agricultural potential. Tamakalte: Black grey soil with a fine clayey texture, resistant to dryness, has a high water retention capacity, and has a good agricultural potential.

126

Region VII: Erfoud (arid and Saharan region) Terrist: White–yellow soil, without stones, no salt, and agricultural potential is low. Amarigh: Different colors of soil (white, yellow, red, and black) and has salinity problems. Azegzou: Schistous nonsoil, agriculture is not practicable. Azougakh: Red soil, may have or may have not stones, and good for agriculture. Amerdoum mouharmel: Very heavy clayey soil, dryness is very slow, and has spontaneous vegetation.

CONCLUSIONS In the Maghreb Arab world, soil science, in general, and soil classification systems, in particular, have followed the same evolution as in all the circum Mediterranean countries. In this region, agricultural practices go back to the prehistoric period, and soil names used before the development of a scientific soil classification scheme were mostly related to local agricultural practices and were based essentially on color, texture, structure, fertility, richness of stones, and water-holding capacity. In Morocco, as likely in the other countries of the Arab world, because of the multiplicity of ethnic tribes and natural conditions (climates, geological material, vegetation), most of these names are maybe the same to qualify different soils.



REFERENCES 1. Finkl, C.W. Soil classification. Benchmark Pap. Soil Sci. 1982, 1, 391 p. 2. Del Villar, E.H. Types de sol de l’Afrique du Nord; Soil Types of North Africa; Fascicule II: Tunis-Rabat, 1947; 137–288. 3. Commission of pedology and cartography of soils (CPCS). In Reprinted from Classification des Sols; INRA, Laboratoire de Geólogie–Pe´dologie: Paris, 1967; pp. 1, 5–13, 15. 4. FAO-UNESCO. Soils of the World; Elsevier Science Publishers B.V., 1987; ELSEVIER=ISRIC. 5. Soil Survey Staff. Soil taxonomy, a basic system of soil classification for making and interpreting soil surveys. In Agriculture Handbook No 436; Soil Conservation Service, U.S. Department of Agriculture, 1975; 754 pp. 6. Atelier sur les bases de donneés SOTER dans les pays de l’UMA; Workshop on SOTER Data Base in UMA Countries; Organisation des Nations Unies pour l’Alimentation et l’Agriculture, Bureau Sous-Re´gional pour l’Afrique du Nord, SNEA: Tunis, 2001; 132 pp. 7. Soil, the Yearbook of Agriculture; United State Department of Agriculture: Washington, DC, 1957; 784 pp. ^ b al Fila ^ ha). 8. Ibn-al-Awam Le livre de l’agriculture (Kita In Traduction from Arab by J.J. Cle´ment-Muller Introduction of Mohammed El Faı¨z, ‘‘Thesaurus’’; Actes Sud=Sindbad, 2000; 1027 pp. 9. Del Villar, E.H. Me´thodes de Classification et Analyse des Sols. Base Scientifique Pour Leur Cartographie Harmonique Universelle; Methods of Classification and Soil Analysis. A Scientific Basis for Their Cartography; 1953; 193 pp. 10. Refleh, Ph.; Chami, A.M. Geography of the Arab World; Annahda Library: Egypt, 1962; 375 pp. Published in Arabic.

Archaeology and Soil John E. Foss University of Tennessee, Knoxville, Tennessee, U.S.A.

INTRODUCTION Geology has had a long period of interaction with archaeology, but the use of soil investigation in archaeology has a rather short history. In 1942, Nikiforoff[1] used the term archeopedology for those soil scientists working with fossil soils or paleosols. Early studies of soils at archaeologic sites were concerned mainly with soil chemical properties (e.g.,[2–4]. An early book by Cromwall[5] also played an important role in demonstrating the usefulness of soil–archaeologic interactions. The past 30 or 40 years have seen a substantial increase in the multidisciplinary effort between these two sciences and have involved more subdisciplines of soil science.

by Scudder, Foss, and Collins[9] Soil Science Society of America Special Pub. No. 44 on ‘‘Pedological Perspectives in Archaeological Research,’’[10] and articles in the Proceedings of Conferences on Pedoarchaeology[11–12] have raised pedologists’ and archaeologists’ awareness of the potential contributions of soil studies to site evaluation. The periodical Geoarchaeology: An International Journal has also been valuable in promoting earth science activity in archaeologic investigations. Some of the major pedologic contributions to field archaeology have included the following:[13]

ROLE OF SOIL SCIENCE IN ARCHAEOLOGY

As archaeologists become more interested in a complete understanding of the chronology and environmental history of sites, a multidisciplinary effort is absolutely necessary. Team members commonly include scientists from soil science, geology, botany, zoology, palynology, and other specializations. Soil science, especially the pedology area (i.e., the study of soil formation and classification), has been particularly active within the past few decades in evaluation of archaeologic sites. Pedology, geology, and other earth sciences often work in the specialized field of geoarchaeology, which means using earth science principles to study archaeologic sites. Fig. 1 shows a landscape of Tikal, Guatemala (Mayan site); this site is one of the many important archaeologic sites that has required the expertise of pedologists to help interpret chronology and land use.[6–9] The study of soils and landscapes is an integral part of many archaeologic investigations. Some federal and state regulations that require geologic and soil input on archaeologic sites have also been responsible for including earth scientists in these studies. Publications in the past decade have indicated the interest of pedologists, geologists, and archaeologists in evaluating soils and landscapes as part of overall archaeologic investigations. Publications such as ‘‘Soils in Archaeology’’ edited by V.T. Holliday,[8] ‘‘Soil Science and Archaeology’’


Determining site delimitation General pedologic stratigraphy Soil–landscape relationships Identification of geologic parent material Correlating soil morphology and archaeologic levels Identifying lithologic (parent material) and pedologic (soil weathering) discontinuities Approximating soil age Identifying paleosols (fossil soils) Contributing to the overall interpretation of site

In the past decade, many of the above pedologic contributions to archaeology were made during the final phase of archaeologic field work. More recently, pedologists have been more involved in phase 1 activity of archaeologic investigations. The early identification of major stratigraphic zones, preliminary analysis of landscape and soil age, and model of site development have resulted in more efficient archaeologic excavations and interpretations.

FIELD STUDIES Archaeologic sites occur in many different geologic provinces and landscape positions. Determination of the site context is thus the most important initial stage in pedoarchaeology. Geologic maps can provide general knowledge of a region, but detailed soil surveys provide the most useful introduction to a study area. These maps produced by the National Resource Conservation Service (NRCS), in cooperation with the Land Grant Institutions, are usually on a county-wide 127

128

Archaeology and Soil

Fig. 1 Landscape view of Mayan city of Tikal, Guatemala.

basis using an air photo base with a scale of 1 : 15840 to 1 : 24000. At this scale, there is not sufficient detail to relate the morphology of individual soil mapping units to specific horizons encountered at an archaeologic site. The landscape and physiographic position of each soil mapping unit, however, are still useful in preliminary analysis of archaeologic sites. The most important and informative archaeologic sites occur in landscapes where sediment is added to a pre-existing surface, therby protecting the artifacts and soil horizons. Those buried sites may occur in the following areas or situations:

Alluvial deposition

Volcanic activity

Eolian deposition Colluvial slopes Mass movement or slumping Seismic areas Artificial deposition or destruction

These situations provide the opportunity for soil burial (subsequently termed paleosols) and archaeologic levels. The buried surfaces (A horizons) of these paleosols are particularly good sources of artifacts and living surfaces when the events above took place in the Holocene. Holliday[8] provides an excellent background in the use of paleopedology in archaeology.

Soil Morphology Soil morphology (e.g., a detailed description of soil profiles) provides the key to understanding and interpreting soils and landscapes at archaeologic sites. The unique soil morphology of a given region and site results from the weathering processes regulated by the interaction of soil-forming factors.[14] These factors are climate, biotic, geology, topography, and length of time that the weathering processes have been operating. The morphologic properties of soils usually described in excavations and their interpretation for archaeologic sites are given in Table.1 A great deal

Table 1 Morphologic properties of soils and their interpretive value at archaeologic sites Soil property

Useful interpretative features

Texture

Lithologic and pedologic discontinuities; classification of geologic materials; determination of argillic horizons; determine relative energy of alluvial sedimentation

Structure

Relative abundance of macropores and potential artifact movement; degree of development is an indicator of soil age; development of clay or organic coatings on argillic horizons (e.g., continuous clay coatings on pedologic faces indicate 10,000 years of development while discontinuous coatings may indicate 4000–5000 years of weathering in southeastern U.S.)

Color

Indicator of organic matter and free iron content; classification of sediments; delineation of horizons; drainage characteristics (redoximorphic features)

Boundary

Abrupt boundary indicator of Ap (plow zone) or recent deposition; boundary becomes more diffuse with age

Consistence

Indicator of structural development, cementation, or consolidation (e.g., recent alluvium usually very friable or loose)

Clay coatings

Coatings on peds or in pores indicate state of development and age

Carbonate

Secondary CaCO3 leaching, coatings, pore filling, and cementation can provide soil age estimates and climatic implications

Horizon identification

Indicates many weathering processes occurring in profile, e.g., A ¼ organic matter accumulation; E ¼ leached zone; Bt ¼ argillic horizon with minimal 4,000 year age; distinguish natural vs. artificial horizons; horizon thickness (solum) is a measure of length of weathering time; diagnostic horizons useful indicator for archaeological interpretation (e.g., argillic, cambic, fragipan, spodic, etc.)

(Modified from Ref.[13].)



of experience and technique is needed to provide an accurate and informative soil description. Evaluation of the age of soils, for example, requires an integration of all the morphologic features that are detailed in Table 1. Certain soil horizons provide general age estimates based on the length of time needed for weathering processes to develop specific features (e.g., argillic horizons). As noted in Table,1 a minimal argillic horizon can form in 4,000 years. Other age estimates of soil horizons have been published previously.[9,15] One of the most useful applications of soil morphology in archaeologic site interpretation is that of distinguishing ‘‘natural’’ from ‘‘artificial’’ or ‘‘man-influenced’’ horizons. Some natural horizons—such as a spodic (Bh) with a dark-colored, organic-rich matrix—may appear as a buried surface or midden. Some albic E horizons could be interpreted as ash layers. Other characteristics that are related to soil genesis, such as redoximorphic features (i.e., mottling or gleying), result from water table fluctuations and often cause confusion in interpretation of color in archaeologic levels. Horizons with calcium carbonate filling (Bk or Ck) have sometimes been identified as plaster-filled.

LABORATORY Laboratory soil characterization for archaeologic interpretations is used to verify and supplement field morphology. Laboratory analysis without complete soil morphology is generally of minimal value for archaeologic interpretation. Complete sampling of all soil horizons, columns, or archaeologic levels is also important to realize the full benefit of the additional cost and labor of soil analysis. Those laboratory analyses that are frequently applied in pedoarchaeology are organic carbon,[16] particle size distribution,[17] and elemental composition.[18] Other soil analysis may include pH, electrical conductivity, mineralogy, free iron, scanning electron microscopy (SEM), energy dispersive x-ray (EDAX), calcium carbonate, and micromorphology. The micromorphologic studies by Goldberg[19] and Macphail and Goldberg [20] have been especially useful in interpreting site stratigraphy and pedologic and geologic events.

FUTURE In the past few decades, pedologists have grown increasingly interested in work on archaeologic sites, and it is likely this trend will continue well into the future. Although we aid archaeologists in understanding the soils and pedologic features they carefully excavate, we have learned a great deal about weathering


129

rates, horizon formation, and landscape development by teamwork with archaeologists and geologists. The use of additional techniques or applications by geoarchaeologists, such as x-ray diffraction (XRD), SEM, EDX, electrical resistivity, ground-penetrating radar, magnetic susceptibility, and micromorphology, will improve soil interpretation work in the future. Despite advances in analytical tools, the key to archaeologic site interpretation still remains the accurate, complete soil morphologic descriptions.

REFERENCES 1. Nikiforoff, C.C. Introduction to paleopedology. Am. J. Sci. 1943, 41, 194–200. 2. Dietz, E.F. Phosphorus accumulation in soil of an Indian habitation site. Am. Antiquity 1957, 22, 405–409. 3. Cook, S.F.; Heizer, R.F. Studies on the chemical analysis of archaeological sites. Univ. Calif. Pub. Anthropol. 1965, 2. 4. Sokoloff, V.P.; Carter, G.F. Time and trace metals in archaeological sites. Science 1952, 116, 1–5. 5. Cromwall, I.W. Soils for the Archeologist; Phoenix House Ltd.: London, 1958. 6. Olson, G.W. Soils and the Environment: A Guide to Soil Surveys and Their Application; Chapman and Hall: New York, 1981. 7. Foss, J.E. Paleosols of pompeii and oplontis. In Stvdia Pompeiana and Classics; Curtis, R.L., Ed.; Orpheus Pub. Inc.: 1988; 127–144. 8. Soils in archaeology; Holliday, V.T., Ed.; Smithsonian Institution Press: Washington, 1992. 9. Scudder, S.J.; Foss, J.E.; Collins, M.E. Soil science and archaeology. Advances in Agronomy 1996, 57, 1–76. 10. Collins, M.E., Carter, B.J., Gladfelter, B.G., Southard, R.J., Eds.; Pedological perspectives in archaeological research. Soil Sci. Soc. Am. Special Pub.; 1995; Vol. 44, 157 pp. 11. Foss, J.E., Timpson, M.E., Morris, M.W., Eds.; Proceedings of the First International Conference on Pedo-Archaeology. Univ. of Tennessee, Agr. Exp. Sta., Special Pub. 1993; Vol. 93–03, 210 pp. 12. Goodyear, A.C., Foss, J.E., Sassaman, K.E., Eds.; Proceedings of the Second International Conference on Pedo-Archaeology. South Carolina Institute of Archaeology and Anthropology, Univ. of South Carolina, Anthro. Studies. 1994; Vol. 10, 157 pp. 13. Foss, J.E.; Lewis, R.J.; Timpson, M.E.; Morris, M.W.; Ammons, J.T. Pedologic approaches to archaeological sites of contrasting environments and ages. Proceedings of the First International Conference on PedoArchaeology; Foss, J.E., Timpson, M.E., Morris, M.W., Eds.; Univ. of Tenn.; Agr. Exp. Sta. Spec. Pub., 1992; Vol. 93–03, 19–22. 14. Jenny, H. Factors of Soil Formation; McGraw-Hill: New York, 1941. 15. Foss, J.E.; Lewis, R.J.; Timpson, M.E. Soils in alluvial sequences: some archaeological implications.

130

In Pedological Perspectives in Archaeological Research; Collins, M.E., Carter, B.J., Gladfelter, B.G., Southard, R.J., Eds.; Soil Sci. Soc. Am. Special Pub., 1995; Vol. 44, 1–14. 16. Stein, J.K. Organic matter in archaeological contexts. In Soils in Archaeology; Holliday, V.T., Ed.; Smithsonian Institute Press, 1992; 193–216. 17. Timpson, M.E.; Foss, J.E. The use of particle-size analysis as a tool in pedological investigations of archaeological sites. Proceedings of the First International Conference on Pedo-Archaeology; Collins, M.E., Carter, B.J., Gladfelter, B.G., Southard, R.J., Eds.; Univ. of Tenn. Agr. Exp. Sta. Spec. Pub., 1992; Vol. 93–03, 69–80.



18. Schuldenrein, J. Geochemistry, phosphate fractionation, and the detection of activity areas at prehistoric North American sites. In Pedologic Perspectives in Archaeological Research; Collins, M.E., Ed.; Soil Sci. Soc. Am. Special Pub., 1995; Vol. 44, 107–132. 19. Goldberg, P. Micromorphology, soils, and archaeological Sites. In Soils in Archaeology; Holliday, V.T., Ed.; Smithsonian Institution Press: Washington, DC, 1992; 45–167. 20. Macphail, R; Goldberg, P. Recent advances in micromorphological interpretation of soils and sediments from Archaeological sites. In Archaeological Sediments and Soils; Barham, A.J., Macphail, R.I., Eds.; Institute of Archaeology, University College: London, 1995.

Arid Soils H. Curtis Monger New Mexico State University, Las Cruces, New Mexico, U.S.A.

INTRODUCTION Scarcity of rain is the dominant characteristic of arid soils. While age, parent material, carbonate, and salt content may vary from arid soil to arid soil, dryness is common to all. Of the total ice-free land area on Earth (130,797,000 km2), about 22% or 28,703,000 km2 is occupied by soils with aridic moisture regimes.[1] Although arid (L. aridus, dry) signifies lack of moisture, technical definitions of arid vary. In some cases, the arid–semiarid boundary is placed at 25 cm (10 in.) of annual rainfall.[2] In other systems, such as the Ko¨ppen–Geiger–Pohl and Meigs systems, the arid (desert)–semiarid (steppe) boundary is based on a combination of rainfall and temperature.[3,4] Still other systems, such as those by Strahler and Soil Survey Staff, use soil moisture to define arid zones because the availability of moisture to plants is more important than annual precipitation itself.[4,5] In all cases, however, rainfall is insufficient to maintain perennial streams. Soils in these regions are unique because relatively little water percolates deep enough to reach groundwater. As a result, carbonates, gypsum, and more soluble salts acumulate in the profiles of many arid soils.

ARID SOILS OF RIVER FLOODPLAINS Floodplain soils along rivers that flow through arid climates were sites of several ancient and eminent civilizations. Sumerian (ca. 3600 B.C.) and later Babylonian (ca. 2000 B.C.) civilizations grew into centers of trade and government as a result of irrigated agriculture on the Tigris and Euphrates River floodplains.[6] Likewise, soils and irrigated agriculture along the Nile of ancient Egypt, the Indus of ancient India, and the Hoang-Ho (Yellow River) of ancient China made it possible for civilizations to create notable schools, calendars, armies, mathematics, medicine, literature, philosophy, science, and art. In the western hemisphere as well, Hohokam, Aztec, and Inca societies emerged in arid and semiarid environments.[7] These civilizations existed because floodplain soils are well suited for irrigated agriculture if groundwater tables are sufficiently deep and salts do not accumulate. In the case of the Nile prior to dam construction, the river would rise and spill over its banks, flood the Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120015640 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

adjacent plain, deposit sediment, and leach salts.[7] In the case of the Tigris and Euphrates, however, drainage canals were needed to carry away leached salts, and with their demise soils became saline.

ARID SOILS ON UPLANDS Most soils in arid regions are not on floodplains, but occur in vast upland areas composed of three major landforms: mountains, piedmont slopes, and basin floors.[8] These major landforms, in turn, are composed of smaller, component landforms. Typically, soil boundaries correspond to component landforms. In mountains, for example, soil boundaries match the boundaries of colluvial wedges, valley fills, and pediments.[9] On piedmont slopes, soil boundaries parallel the boundaries of alluvial fans, ballenas, and fan skirts. On basin floors, which characteristically have little topographic relief, soil boundaries generally follow landforms produced by wind, such as deflational blowouts, dunes, and eolian plains, or landforms produced by pluvial lakes, such as lake plains, playas, and beach plains. Of the five soil-forming factors (climate, time, biota, topography, and parent material), climate is the defining factor of arid soils, although time is an important factor as well. The impact of time on arid soils is revealed by carbonate and clay accumulations in soils of progressively older geomorphic surfaces (Fig. 1). Carbonate in nongravelly soils, for example, progresses from carbonate filaments in middle Holocene soils to carbonate nodules in late Pleistocene soils to carbonate-indurated horizons in middle Pleistocene soils.[10] Clay likewise accumulates with time to form argillic horizons. However, the correlation of clay accumulation with time is less robust than carbonate accumulation with time because many ancient soils that have calcretes do not have argillic horizons.[11] This indicates that argillic horizons are more vulnerable to obliteration by erosion and bioturbation than calcic or petrocalcic horizons. Arid soils are not only unique because carbonate, gypsum, and soluble salt accumulate, but also because many have vesicular A-horizons covered by desert pavement[12] or microbiotic crust.[13] In addition, inadequate water and nitrogen suppress biomass 131

132

Arid Soils

Fig. 1 Desert piedmont slope rising to a mountain chain in southern New Mexico. Progressively older geomorphic surfaces with progressively greater soil development are typical features of piedmont slopes. The younger soil on the right (about 3000 yr old) has a small amount of carbonate (white zone in profile). The older soil to the left (25,000–150,000 yr old) has substantially more carbonate. In addition, the older soil has an argillic horizon overlying the carbonate horizon.

production on arid soils to about one-tenth the biomass of temperate forest soils.[14] Nevertheless, soil animals such as rodents, ants, and termites are common. Ants, for example, can transfer 80 g=m2 of desert soil to the land surface per year, which is as much as ants transfer in more mesic environments.[15]

TYPES OF ARID SOILS The main criterion for the classification of arid soils is soil dryness, or the aridic (torric) moisture regime, which is defined as soils too dry for agricultural crops unless irrigated.[16] Further taxonomic subdivisions are based on diagnostic horizons. In contrast to the notion that arid soils are poorly developed, as written in some soil science books, many arid soils are strongly developed with a variety of diagnostic subsurface horizons.[17] These horizons include the argillic, natric,


salic, gypsic, petrogypsic, calcic, and petrocalcic horizons, and the duripan.[5] Diagnostic surface horizons include the ochric epipedon with minor occurrences of the mollic and anthropic epidons. Arid soils that have diagnostic subsurface horizons are generally classified as Aridisols. These include many of the older soils on piedmont slopes, basin floors, mountain uplands. Various types of Aridisols occur on the landscape because of lateral changes in particle size, truncation of diagnostic horizons, degradation of diagnostic horizons, moisture heterogeneity across the landscape, and age differences, which can range from Historical to Pliocene within small geographical areas.[9,18] Arid soils that lack diagnostic subsurface horizons are generally classified as Entisols, which fall into the azonal concept of Sibirtsev.[19] These include many of the younger soils on floodplains, dunes, and erosional surfaces. In Soil Taxonomy, floodplain soils are mainly

Arid Soils

classified as Fluvents or, more specifically, Torrifluvents.[5] Arid soils associated with dunes are commonly Torripsamments and those associated with erosional surfaces are commonly Torriorthents. and Entisols Aridisols (14,942,000 km2) 2 (12,682,000 km ) are the dominant soil types in arid regions, although other soil types include Vertisols (889,000 km2) and Oxisols (31,000 km2) and very minor amounts of Mollisols, Andisols, Histosols, and Spodosols.[1] Arid soils grade into semiarid soils across three climatic transects: laterally into wetter regions, upslope into wetter climates at higher elevations, or downslope into run-in areas with wetter microclimates. Taxonomically, changes in soil types from dry region aridic to wetter region ustic or xeric moisture regimes are expressed at the Suborder and Great Group level (Fig. 2). Linked to this climatic transition is a progressive change in vegetation—desert shrublands give way to grasslands that in turn give way to wood-lands. Also across this transition, soils have progressively deeper carbonate horizons. In the Chihuahuan Desert, for instance, carbonate zones are 50 cm deep at 230 mm of annual rainfall and 100 cm deep at 320 mm of annual rainfall.[20] Likewise, gypsum zones progressively deepen from about 50 cm depth at 150 mm of annual rainfall to about 100 cm depth at 250 mm annual rainfall.[21] Accompanying an increase in rainfall is an increase in soil organic matter. Although the amount of organic matter depends on the clay content, organic matter ordinarily increases from less than 0.5% in A-horizons of arid shrubland soils to 2–5% in semiarid and subhumid grassland soils.[22]

133

ECOLOGICAL SIGNIFICANCE Biodiversity in arid regions is linked to habitat diversity. Habitat diversity, in turn, is created by various microclimates caused by topographic factors and soil properties.[23] Thus, soils help mold and are molded by ecosystems. In many arid regions, such as the southwestern United States, soils have been impacted by ecosystem changes of the Holocene and late Pleistocene when wetter climates alternated with drier climates.[24,25,26,27] According to this model, landscape stability was greater during wetter climates because denser vegetative cover reduced erosion. With reduced erosion, soil formation occurred. In contrast, instability was greater during drier climates because sparse vegetative cover gave rise to more bare ground and increased erosion. With increased erosion, soil formation was inhibited. This oscillation between stability and instability is recorded as stacked sequences of buried paleosols in depositional environments and as stepped sequences of fan-terraces in areas that grade to fluctuating river base-levels. Globally, arid soils affect atmospheric dust, rain chemistry, ocean fertilization, albedo, denitrification, and the carbon cycle as both sinks and sources of CO2.[28,29] Carbonate–carbon, for instance, is the second largest terrestrial carbon pool, totaling approximately 50–60 Pg C in the dryland zones of the U.S.[30] and approximately 750–950 Pg C in the dryland zones of the world.[31] Humans have lived on arid soils for millennia. In fact, the oldest known hominid tools are in arid East Africa and date back 2.5 million yr.[32] Today arid

Fig. 2 Illustration of Soil Taxonomy Suborders and Great Groups that have aridic moisture regimes (shaded) and their moister counterparts that have ustic and xeric moisture regimes. (From Ref.[5].)


134

soils are important to humans for livestock grazing, irrigated agriculture, and urban development. In many arid regions of the world, human land use has resulted in desertification and diminishing groundwater supplies, both of which are increasingly important social and scientific issues as human population increases. ACKNOWLEDGMENTS Grateful acknowledgment is made to Haiyang Xing and Marco Inzunza for making the figures and Rebecca Kraimer for reviewing the manuscript. REFERENCES 1. Wilding, L.P. Introduction: general characteristics of soil orders and global distributions. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, FL, 2000; E175–E182. 2. Bull, W.B. Geomorphic Responses to Climatic Change; Oxford University Press: New York, 1991; 326 pp. 3. Dick-Peddie, W.A. Semiarid and arid lands: a worldwide scope. In Semiarid Lands and Deserts; Skuji}s, J., Ed.; Marcel Dekker: New York, 1991; 3–32. 4. Strahler, A.N.; Strahler, A.H. Modern Physical Geography, 3rd Ed.; Wiley: New York, 1987; 544 pp. 5. Soil survey staff. Soil Taxonomy—A Basic System of Soil Classification for Making and Interpreting Soil Surveys, 2nd Ed.; USDA Agriculture Handbook Number 436; U.S. Govt. Printing Office: Washington, DC, 1999. 6. Durant, W. Our Oriental Heritage; Simon and Schuster: New York, 1935; 1049 pp. 7. Dregne, H.E. Soils of Arid Regions; Elsevier: Amsterdam, 1976; 237 pp. 8. Peterson, F.F. Landforms of the Basin and Range Province Defined for Soil Survey; Nevada Agricultural Experiment Station, Tech. Bull. 28, Univ. of Nevada: Reno, 1981; 52 pp. 9. Gile, L.H.; Hawley, J.W.; Grossman, R.B. Soils and Geomorphology in the Basin and Range Area of Southern New Mexico—Guidebook to the Desert Project; New Mexico Bureau of Mines and Mineral Resources, Memoir 39, Socorro: New Mexico, 1981; 222 pp. 10. Gile, L.H.; Peterson, F.F.; Grossman, R.B. Morphology and genetic sequences of carbonate accumulation in desert soils. Soil Sci. 1966, 101, 347–360. 11. Gile, L.H. Eolian and associated pedogenic features of the jornada basin floor, southern new mexico. Soil Sci. Soc. Am. J. 1999, 63, 151–163. 12. McFadden, L.D.; Wells, S.G.; Jercinovich, M.J. Influences of eolian and pedogenic processes on the origin and evolution of desert pavements. Geology. 1987, 15, 504–508. 13. Kidron, G.J.; Yaalon, D.H.; Vonshak, A. Two causes for runoff initiation on microbiotic crusts: hydrophobia and pore clogging. Soil Sci. 1999, 164, 18–27. 14. Ludwig, J.A. Primary productivity in arid lands: myths and realities. J. Arid Environ. 1987, 13, 1–7.


Arid Soils

15. Whitford, W.G.; Schaefer, D.; Wisdom, W. Soil movement by desert ants. The Southwestern Naturalist. 1986, 31, 273–274. 16. Smith, G.D. The Guy Smith Interviews: Rationale for Concepts in Soil Taxonomy; Soil Management Support Service Monograph No. 11; U.S. Government Printing Office: Washington, 1986. 17. Ahrens, R.J.; Eswaran, H. The international committee on aridisols: deliberations and rationale. Soil Surv. Horizons. 2000, 41, 110–117. 18. Gile, L.H. Causes of soil boundaries in an arid region; I. age and parent materials. Soil Sci. Soc. Am. Proc. 1975, 39, 316–323. 19. Sibirtsev, N.M. Selected Works, Soil Science; Issued in Translation by the Israel Program for Scientific Translation: Jerusalem, 1966; Vol. 1, 354 pp. 20. Gile, L.H. Holocene soils and soil-geomorphic relations in a semi-arid region of southern new mexico. Quatern. Res. 1977, 7, 112–132. 21. Cooke, R.; Warren, A.; Goudie, A. Desert Geomorphology; UCL Press: London, 1993; 526 pp. 22. Birkeland, P.W. Soils and Geomorphology, 3rd Ed.; Oxford University Press: New York, 1999; 430 pp. 23. McAuliffe, J.R. Landscape evolution, soil formation, and ecological patterns and processes in sonoran desert bajadas. Ecol. Monogr. 1994, 64, 111–148. 24. Ruhe, R.V. Age of the rio grande valley in southern New Mexico. J. Geol. 1962, 70, 151–167. 25. Gile, L.H.; Hawley, J.W. Periodic sedimentation and soil formation on an alluvial-piedmont in southern New Mexico. Soil Sci. Soc. Am. Proc. 1966, 30, 261–268. 26. Monger, H.C.; Cole, D.R.; Gish, J.W.; Giordano, T.H. Stable carbon and oxygen isotopes in quaternary soil carbonates as indicators of ecogeomorphic changes in the northern chihuahuan desert, USA. Geoderma 1998, 137–172. 27. Buck, B.J.; Monger, H.C. Stable isotopes and soilgeomorphology as indicators of holocene climate change, northern chihuahuan desert. J. Arid Environ. 1999, 43, 357–373. 28. Schlesinger, W.H.; Reynolds, J.F.; Cunningham, G.L.; Huenneke, L.F.; Jarrell, W.M.; Virginia, R.A.; Whitford, W.G. Biological feedbacks in global desertification. Science 1990, 247, 1043–1048. 29. Schlesinger, W.H. Biogeochemistry: An Analysis of Global Change, 2nd Ed.; Academic Press: New York, 1997. 30. Monger, H.C.; Martinez-Rios, J.J. Inorganic carbon sequestration in grazing lands. In The Potential of U.S. Grazing Lands to Sequester Carbon and Mitigate the Greenhouse Effect; Follette, R.F., Kimble, J.M., Lal, R., Eds.; CRC Press: Boca Raton, FL, 2001; 87–118. 31. Eswaran, H.; Reich, P.F.; Kimble, J.M.; Beiroth, F.H.; Radmanabhan, E.; Moncharoen, P. Global carbon sinks. In Global Climate Change and Pedogenic Carbonates; Lal, R., Ed.; CRC Press: Boca Raton, FL, 2000; 15–26. 32. Ambrose, S.H. Paleolithic technology and human evolution. Science 2001, 291, 1748–1753.

Atterberg Limits Thomas Baumgartl Institute for Plant Nutrition and Soil Science, Kiel, Germany

INTRODUCTION Soil is exposed to different states of stability depending on the amount of water that it contains. This characteristic is described as consistency (refer to entry on soil consistency and plasticity) and specifies the state of a remolded and cohesive soil in the range from the liquid (when wet) to plastic and finally solid (when dry) state. Different soils contain a specific amount of water at these different states of stability. In 1911, the Swedish soil physicist Atterberg developed a classification system and method with which these states of consistency could be determined. The method is based on the determination of the water content [calculated as: (mass of water)=(dry mass of soil)] at distinct transitions between different states of consistency of soil. These transitions are defined as liquid limit, plastic limit and shrinkage limit, and are generally referred to as Atterberg limits. The values for these limits are dependent on various soil parameters, e.g., particle size, specific surface area of the particles which is able to take up water, and hence its particle size distribution. These limits are used to derive indices, e.g., index of plasticity and index of consistency, and are often used for the mechanical characterization of soils.

DEFINITION OF LIMITS AND THEIR DETERMINATION Liquid Limit (Upper Plastic Limit) wL The liquid limit describes the transition from a viscous liquid to a plastic state. Soils with a water content at the liquid limit barely flow under an applied force. The associated capillary forces of the water menisci in the unsaturated pore system are equivalent to pF 0.5 (0.3 kPa matric potential).[1] The liquid limit is determined by a method and device developed by Casagrande. The principle is to find the water content (kg=kg) at which a soil sample starts to liquify under a small applied stress. In practice a groove is cut into soil samples with different water contents. These soil samples are then exposed to a small standardized force by repeatedly dropping the Casagrande cup over a distance of 10 mm until the groove is close to ca. 10 mm. A semilogarithmic plot of the number of Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001729 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

blows as a function of water content will result in a straight line. The liquid limit is defined as the water content at 25 blows (Fig. 1). Plastic Limit (Lower Plastic Limit) wP The plastic limit determines the transition from a plastic (cohesive) to a semirigid or brittle state. Under an applied force cracking will occur. In sandy soils the plastic limit often cannot be determined. The matric potential at the plastic limit is the main cohesive stress and ranges between 63 and 200 kPa (pF 2.8–3.3).[1] The plastic limit is determined by forming a moist ball from 2 to 3 g of soil, which is then rolled on a piece of frosted glass to a rod of thickness ca. 3 mm. The remolding and rolling is repeated until the 3 mm rod starts to break up into pieces of 10–20 mm. The gravimetric water content (kg=kg) at this point gives the plastic limit. Shrinkage Limit wS Cohesive and remolded soils reduce their soil volume with the loss of water due to capillary forces. If in a drying process the reduction of the total soil volume equals the volume of water loss, then the soil shows normal shrinkage. Below a certain water content, the further shrinkage of the soil volume is restricted due to a high number of particle contact points and high effective stresses. This restricted shrinkage pattern is called residual shrinkage (Fig. 2). The transition from normal to residual shrinkage defines the shrinkage limit. The shrinkage limit is often calculated by wS ¼ 0:65 wP Index of Plasticity IP The index of plasticity is the amount of water between plastic and liquid limit and is calculated by IP ¼ wL wP It describes the sensitivity in the mechanical behavior of a soil towards changes in water content. However, it does not explain mechanical stability as hydraulic parameters are not included which are necessary as water flow becomes important when stresses are applied on a 135

136

Atterberg Limits

The water content w as such provides no information about the consistency of soils. The same water content in a sandy soil may reflect a liquid state, where as in a clay soil the behavior may be brittle. The index of consistency normalizes the water contents and characterizes whether the soil is close to the plastic limit ðmax IC ¼ 1Þ or the liquid limit ðmin IC ¼ 0Þ.

Number of blows

100

Index of Shrinkage ISch

25

The difference between the plastic limit and the shrinkage limit results in the index of shrinkage: ISch ¼ wP wS 10 0.2

0.3

0.4

0.5

Water content

wL

Soils with water contents within this range are most suitable for cultivation (see Fig. 2). Table 1 summarizes some general values for the Atterberg limits and indices.[8]

Fig. 1 Determination of the liquid limit.

soil sample. Values for the index may range from 0 (no plastic behavior) for sandy material to 1 (100%) for clay. Index of Consistency IC The index of consistency is determined as the ratio of the difference between the liquid limit and actual water content and the index of plasticity: IC ¼

wL w wL w ¼ wL wP IP

Specific volume [cm3/g soil]

FACTORS INFLUENCING THE ATTERBERG LIMITS Many mechanical processes are linked to hydrological properties of soils. Therefore, values of the limits and indices are influenced by factors which are generally important for the water retention curve, e.g., the capacity of swelling and shrinkage, clay content, type of clay minerals and organic matter. Generally, the values of the limits and indices increase with their clay content. As the liquid limit increases in comparison to the plastic limit, the index of plasticity also increases. The swelling and shrinkage intensity is dependent not only on the amount but also on the type of clay mineral. Skempton introduced a factor described as the activity of clay:[2,9] A ¼

Normal shrinkage

IP : % clay content

Shrinkage

The values of A can be classified as: 1) A > 1.25: active soil with high capacity of swelling and shrinking [Ca-montmorillonite (A 1.5), Na-montmorillonite Table 1 Consistency limits (g water=g soil) for different soil textures

Residual shrinkage

Texture ISch

wS

Ip

wP wL Water content [g/g]

Fig. 2 Shrinkage behavior with change in water content; relations to Atterberg limits wS, wP, wL and Atterberg indices. (From Ref.[9].)


Sand

Silt

Clay

Liquid limit

0.15–0.20

0.30–0.40

0.40–1.50

Plastic limit

0

0.20–0.25

0.25–0.5

Index of plasticity

0

0.10–0.15

0.10–1.00

0.12–0.18

0.14–0.25

0.08–0.25

Consistency limits

Shrinkage limit (From Ref.[8].)

Atterberg Limits

137

Table 2 Mechanical properties of soils and Atterberg limits Index of consistency Symbol

0

0.25

1

1.3

wP

wS

Index of plasticity Slurry

Unconfined compression strength (kPa) Cultivation (pos., þ; neg., )

0.75

wL

Index Description

0.5

Index of shrinkage

Very soft

Soft

Deformable

Stiff

Medium hard

Hard

400

þ

(From Refs.[2,9].)

(A 7.5), smectite, salt influenced clays]); 2) 0.75 < A < 1.25: normal soils (illite); 3) A < 0.75: inactive clay with only little swelling–shrinking activity (kaolinite). Organic matter increases both the plastic and liquid limits, but does not have a big effect on the index of plasticity. Organic substances in a soil matrix seem therefore to increase the surface hydration. Once this pool for water uptake is saturated, the soil shows the same mechanical behavior towards changes in water content as when organic matter is absent, only at a higher level of water content. Thus, the index of consistency is higher.[1] With the exception of organic matter and clays, the amount and type of exchangeable cations have a significant effect on the value of the limits.[3,5,12] Nasaturated soils reduce the liquid limit, but increase the shrinkage limit. Soils therefore have the tendency to show crust formation at an earlier stage and will slake at lower water contents.[1,8]

and 1. Drier soils increase the energy input needed for cultivation, which can be a serious problem for clay-rich soils as plowing can become difficult. In the case of lower than optimal plasticity indices, the soil structure can be destroyed easily when the soil is kneaded by trafficking resulting in ecological problems. As a result, the hydraulic conductivity and gas flow as well as nutrient uptake of plants can be reduced. Hence, cultivation at index of consistency smaller than 0.75 can have a serious effect on plant growth and soil biological activity. Although the values of the limits and indices are not independent values, they can be related to each other (e.g., IP and wL). Classifications with respect to particle size distribution, geological origin of material, and suitability under soil mechanical point of view can be derived thereafter. With the ratio of liquid limit and index of plasticity a linear relationship was found by Casagrande and described as A-line,[13] following the equation: IP ¼ 0:73ðwL 0:2Þ

APPLICATION


70 increase of plasticity

60

e

lin

A-

Index of plasticity

Soil charcteristics are inherent in the values of the Atterberg limits. Therefore, Atterberg limits are correlated with soil properties. For specific soils, investigations have been carried out which correlate the total particle surface with the plastic limit.[4,6] Soil strength can be defined by its compressibility and compressibility is correlated to the Atterberg limits.[11] As soil strength is influenced to a great extent by the energy status of the capillary water, Atterberg limits reflect the soil water potential and show the dependency on texture and the water retention curve.[10] At a broader scale, the Atterberg limits can be used for the evaluation of the trafficability and cultivation of soils. Table 2 lists the limits and derived mechanical properties and qualities of soil substrates. From this classification (Table 2), it is evident that an optimal range of water contents for agricultural use can be determined. This range is present when the soil is stiff and has a compression strength of >100 kPa and an index of consistency between 0.75

50 40 30

ic

ic

or

d

in

ills

c

10

cohesionless soils

g

i an

or

in

cl

an

n ga

20

s

ay

ys

a cl

g or

an

s

increase of compactiblity

0 0

20 30

40 50 60

80

100

Liquid limit Fig. 3 Classification of soils according to Atterberg limits and Casagrande A-line.

138

Atterberg Limits

Table 3 Mechanical parameters (angle of internal friction, cohesion) dependent on texture and index of plasticity IP

Angle of internal friction }0 ( )

Cohesion (kPa)

Clay of high plasticity: wL > 0.5

0.50–0.75 0.75–1.00 1.00–1.30

17.5 17.5 17.5

0 10 25

Clay and silt of medium plasticity: 0.35 < wL < 0.5

0.50–0.75 0.75–1.00 1.00–1.30

22.5 22.5 22.5

0 5 10

Clay and silt of low plasticity: wL < 0.35

0.50–0.75 0.75–1.00 1.00–1.30

27.5 27.5 27.5

0 2 5

Texture

(Adapted from Ref.[9].)

It distinguishes soil with content of organic matter of selenate. For the anions, the greater the chemical attraction for the surface, the more marked the continuing reaction. The continuing reaction is caused by diffusion of adsorbed

BEHAVIOR OF SOME IONS Before considering the ions in detail, note that selenate, sulfate, phosphate, selenite, and arsenate all form bidentate, inner-sphere complexes with the surface. That is, two of the oxygen atoms provide direct chemical links to the surface atoms. When a reactant forms a bidentate link to the surface, it is appropriate to refer to the divalent ion in solution.[1] Similarly, in the case of a monodentate link, as with borate, it is appropriate to refer to the monovalent ion. Boric acid is fairly weak, with pH at about 9. Therefore, in the normal range of soil pH, the proportion of monovalent borate ion increases 10-fold for each unit increase in pH. The effects of pH on the charge and potential and the effects on acid dissociation therefore oppose each other: the increasingly negative electric potential favors decreased desorption; the increasing dissociation favors increased adsorption. Because the ion is monovalent, the effects of surface charge are not quite large enough to exceed the effects of the increasing value of the dissociation term. Thus, sorption increases with increasing pH. Selenious acid is a diprotic acid with pK1 at 2.7 and pK2 at 8.5 in very dilute solution. The main species present in the range of soil pH values are HSeO3 and SeO32, with the divalent ion increasing with increasing pH. However, because the relevant ion is divalent, the effects of the increasingly negative


Fig. 3 The effect of period of reaction at 25 C for selenite and zinc on the relationship between solution concentration and amount of sorption. The lines were obtained by fitting the model described in the text to the data. (From Ref.[1].)

Chemisorption

ions into the absorbing particle. This is a slow process but can be accelerated by raising the temperature. Consequently, increased temperature increases sorption and/or decreases solution concentration.

DESORPTION The more marked the continuing reaction, the more deeply buried the ion becomes. Although desorption


221

curves do not seem to follow the same track as sorption curves, they can be fully predicted if the continuing reaction is taken into consideration.

REFERENCE 1. Barrow, N.J. The four laws of soil chemistry: the leeper lecture 1998. Aust. J. Soil Res. 1999, 37, 787–829.

Chlorine Uzi Kafkafi The Hebrew University of Jerusalem, Rehovot, Israel

Guohua Xu College of Resources and Environmental Sciences, Nanjing Agricultural University, Nanjing, China

INTRODUCTION Chlorine is found in nature in the form of negatively charged (Cl), highly water-soluble anion. Most of the world’s Cl is found either in oceans or in salt deposits left by evaporation of old inland seas, which are now found in deep quarries. Field and glasshouse studies in the mid-1800s to the early 1900s have shown that the influence of Cl on plant growth varies with the plant variety.[1] Lipman[2] demonstrated the beneficial effect of Cl on buckwheat (Triticum asetivum L.) growth. Arnon and Whatley[3] suggested that Cl is an essential cofactor in the O2 evolution during photosynthesis. The first complete definition of Cl, as a plant essential micronutrient was described by Broyer et al.[4] The function of Cl in crop yield has been largely neglected,[5] as it becomes a limiting factor for plant growth only in areas of high precipitation far from the sea. The negative effects of high Cl concentrations in soil and irrigation water on crop production are observed in coastal, arid and semi-arid areas, where freshwater sources are often scarce and the available groundwater is saline. The dependence of modern agriculture on irrigation and fertilization causes more concern on excess of Cl rather than on its deficiency.[6]

CHLORIDE IN SOIL The concentrations of Cl in natural sources are listed in Table 1. The Cl content in the soil is not an intrinsic property of the soil but rather a result of soil management, rainfall and evaporation, and irrigation and fertilization. The Cl concentration in rainwater ranges from about 20–50 g=m3 close to seashore to 2–6 g=m3 in inner continental areas.[5,7] The annual amount of Cl deposition on land ranges from 17–175 kg=ha. Mid-continental areas such as the Great Plains of North America receive 6.0–7.0

Asteraceae

Lettuce

Lactuca sativa L.

Leaves

>0.14

2.8–19.8

>15.0

References [21,23] [19]

>23.0

[24]

Chenopodiaceae

Spinach

Spinacia oleracea L.

Shoot

>0.13

[25]

Chenopodiaceae

Sugarbeet

Beta vulgaris L.

Leaves

0.71–1.78

[26]

Chenopodiaceae

Sugarbeet

Beta vulgaris L.

Petioles

7.1–7.2

>50.8

[6,26]

Fabaceae

Alfalfa

Medicago sativa L.

Shoot

0.65

0.9–2.7

6.1

[27,28]

Fabaceae

Peanut

Arachis hypogaea L.

Shoot

4.6

0.3–1.5

16.7–24.3

Fabaceae

Red clover

Trifolium pratense L.

Shoot

Fabaceae

Soybean

Glycine max L. Merr.

Leaves

Fabaceae

Subterranean clover

Trifolium subterraneum L.

Shoot

>1.0

Gramineae

Barley

Hordeum vulgare L.

Heading shoot

1.2–4.0

Gramineae

Corn

Zea mays L.

Ear leaves

Gramineae

Corn

Zea mays L.

Shoots

0.05–0.11

Gramineae

Rice

Oryza sativa L.

Shoot

5.3

[29,35]

Rutaceae

Citrus

Citrus sp. L.

Leaves

2.0

4.0–7.0

Solanaceae

Potato

Solanum tuberosum L.

Mature shoot

4.0 1.1–10.0

[5,9] >32.7

[33] [24]

>7.0–8.0

[34]

5.1–10.0

>13.6

[6]

1.5

3.7–4.7

>7.0

[5,29]

1.2–4.0

>4.0

[5,9]

1.5–4.0

7.0

10.0–25.0

>25.0–33.1

[6]

>2.1

[28]

0.1

1.2–10.0 0.7–8.0

[28,35]

[20]

b

10–40 0.25

[20]

>10.0

[6,28]

30.0

[4,38]

10.0–11.0

[6,28] Chlorine

a

0.15–0.21

[29]

Chlorine

cells, and consequent stomata opening. The relative contribution of Cl and malate may vary among species, and depend on the availability of external Cl and plant growth environment.[17] In plant species such as onion (Allium cepa L.), which lack the functional chloroplasts for malate synthesis in the guard cells, Cl is essential for stomata functioning.[17,18] Members of the Palmaceae, such as coconut and oil palm, which might possess starch-containing chloroplasts in their guard cells, also require Cl for stomata function.[19]

INTERACTION OF CHLORIDE UPTAKE WITH THE UPTAKE OF OTHER NUTRIENTS Ammonium is taken up by plant as a cation, and therefore relatively more anions have to be taken up to maintain the electrical neutrality of the uptake process. Plants fertilized with NH4þ, usually contain higher Cl levels in the tissue than plants fertilized with NO3 or with both N sources, irrespective of the Cl concentration in the nutrient solution.[6] Thus, when Cl is present in the root medium, NH4þ uptake may increase the salt sensitivity of the plants. The antagonism between NO3 and Cl uptake has been well demonstrated in many crops.[6] When both Cl and NO3 anions are taken up by the root against their electrochemical gradient, Cl maintains its negative charge, while NO3 is metaboliszed and loose its negative charge. The accumulation of Cl reduces its further uptake since the Cl electrochemical potential gradient builds up during its accumulation in the cell.[20] Nitrate can prevent Cl toxicity of avocado at a concentration of up to 16 mM in the root medium.[20] On the other hand, Cl application may also be used as a strategy to decrease the NO3 content of leafy vegetables, as spinach (Spinacia oleracea L.), lettuce, and cabbage (Brassica oleracea L.), which are classified as NO3 accumulators.[6] Increasing concentrations of Cl in the root medium has generally no consistent effect on K concentrations in the plants. Potassium concentrations in the leaves of kiwifruit are significantly higher for vines receiving KCl than for vines receiving K2SO4 as the kiwifruit uses Cl rather than organic anions for charge balance, and thus maintains a high K uptake.[21]

CHLORIDE MANAGEMENT IN IRRIGATION AND FERTILIZATION Large amounts of Cl enter the field through irrigation. The amount of Cl added to a field with 500 mm of irrigation water containing only 200 g Cl=m3 is only 1000 kg=ha. This is four times more than the amount of Cl applied by fertilization with KCl at 500 kg=ha.


225

The irrigation system influences the distribution of Cl salts in the soil. Irrigation with saline water is managed by an excess of irrigation to meet the leaching requirement for avoiding salt accumulation in the root zone.[22] The amount of water required to wash salts out of the root zone can be estimated from the electrical conductivity (EC) of the irrigation water and the mean EC of the saturated soil extract at which no crop yield reduction occurs under sprinkler or surface irrigation.[22] Fertilization with KNO3 under saline conditions might reduce the toxic effect of salinity of some woody plants even if it is associated with increased soil osmotic potential.[20] Under high salinity and Cl conditions, KNO3 or K2SO4 as K fertilizers are preferred due to their lower salt index and absence of Cl. Addition of adequate P can also be helpful in alleviating salt stress.[6]

CONCLUSIONS Chloride ion is essential for plant growth as a micronutrient. It is taken up by plants in large quantities when its concentration in the soil is elevated due to irrigation, fertilization, and evaporation of water from the soil surface. Cl deficiency was monitored in inland regions with ample rainfall. There is a wide range of Cl concentrations needed by plants, [from Cl-sensitive plants like avocado to tolerant plants like coconuts.] The main load of Cl is observed when irrigation with saline water is employed. The main Cl rich fertilizer is KCl that is normally applied without deleterious effects.

REFERENCES 1. Tottingham, W.E. A preliminary study of the influence of chlorides on the growth of certain agricultural plants. J. Am. Soc. Agron. 1919, 11, 1–32. 2. Lipman, C.B. Importance of silicon, aluminum, and chlorine for higher plants. Soil Sci. 1938, 45, 189–198. 3. Arnon, D.L.; Whatley, F.R. Is chloride a coenzyme of photosynthesis? Science 1949, 110, 554–556. 4. Broyer, T.C.; Carlton, A.B.; Johnson, C.M.; Stout, P.R. Chloride—a micronutrient element for higher plants. Plant Physiol. 1954, 29, 526–532. 5. Fixen, P.E. Crop responses to chloride. Adv. Agron. 1993, 50, 107–150. 6. Xu, G.H.; Magen, H.; Tarchitzky, J.; Kafkafi, U. Advances in chloride nutrition of plants. Adv. Agron. 2000, 68, 97–150. 7. Yaalon, D.H. The origin and accumulation of salts in groundwater and in soils in Israel. Bull. Res. Counc. Israel 1963, 11G, 105–131. 8. Wang, J.H.; Yu, T.R. Release of hydroxyl ions during specific adsorption of chloride by variable-charge soils. Z Pflanzenernahr Bodenk 1998, 161, 109–113.

226

9. Engel, R.E.; Bruckner, P.L.; Mathre, D.E.; Brumfield, S.K.Z. A chloride deficient leaf spot syndrome of wheat. Soil Sci. Soc. Am. J. 1997, 61, 176–184. 10. Maas, E.V. Physiological responses to chloride. In Special Bulletin on Chloride and Crop Production, No. 2; Jackson, T.L., Ed.; Potash & Phosphate Institute: Georgia, 1996; 4–20. 11. Felle, H.H. The Hþ=Cl symporter in root-hair cells of Sinapsis alba. An electrophysiological study using ionselective microelectrodes. Plant Physiol. 1994, 106, 1131–1136. 12. Glass, A.D.M.; Siddiqi, M.Y. Nitrate inhibition of chloride influx in barley: implications for a proposed chloride homeostat. J. Exp. Bot. 1985, 36, 557–566. 13. Rognes, S.E. Anion regulation of lupin asparagine synthetase: chloride activation of the glutamine-utilizing reaction. Phytochemistry 1980, 19, 2287–2293. 14. Churchill, K.A.; Sze, H. Anion-sensitive, Hþ-pumping ATPase of Oat roots. Plant Physiol. 1984, 76, 490–497. 15. Hofinger, M.; Bottger, M. Identification by GC-MS of 4-chloroindolylacetic acid and its methyl ester in immature Vicia faba broad bean seeds. Phytochemistry 1979, 18, 653–654. 16. Flowers, T.J. Chloride as a nutrient and as an osmoticum. Adv. Plant Nutr. 1988, 3, 55–78. 17. Talbott, L.D.; Zeiger, E. Central roles for potassium and sucrose in guard-cell osmoregulation. Plant Physiol. 1996, 111, 1051–1057. 18. Schnabl, H.; Raschke, K. Potassium chloride as stomatal osmoticum in Allium cepa L. (onion), a species devoid of starch in guard cells. Plant Physiol. 1980, 65, 88–93. 19. von Uexkull, H.R. Chloride in the nutrition of coconut and oil palm. Trans. Int. Congr. Soil Sci. 1990, IV (14), 134–139. 20. Bar, Y.; Apelbaum, A.; Kafkafi, U.; Goren, R. Relationship between chloride and nitrate and its effect on growth and mineral composition of avocado and citrus plants. J. Plant Nutr. 1997, 20, 715–731. 21. Buwalda, J.G.; Smith, G.S. Influence of anions on the potassium status and productivity of kiwifruit (Actinidia deliciosa) vines. Plant Soil 1991, 133, 209–218. 22. Keller, J.; Bliesner, R.D. Trickle irrigation planning factor. In Sprinkler and Trickle Irrigation; Keller, J., Bliesner, L.R.D., Eds.; Van Nostrand Reinhold: New York, 1990; 453–477. 23. Prasad, M.; Burge, G.K.; Spiers, T.M.; Fietje, G. Chloride induced leaf breakdown in kiwifruit. J. Plant Nutr. 1993, 16, 999–1012.


Chlorine

24. Johnson, C.M.; Stout, P.R.; Broyer, T.C.; Carlton, A.B. Comparative chloride requirements of different plant species. Plant Soil 1957, 8, 337–353. 25. Robinson, S.P.; Downton, W.J.S. K, Na and Cl content of isolated intact chloroplasts in relation to ionic compartmentation in leaves. Arch. Biochem. Biophys. 1984, 228, 197–206. 26. Ulrich, A.; Ohki, K. Chloride, bromine, and sodium as nutrients for sugar beet plants. Plant Physiol. 1956, 31, 171–181. 27. Ozanne, P.G.; Woolley, J.T.; Broyer, T.C. Chloride and bromine in the nutrition of higher plants. Aust. J. Biol. Sci. 1957, 10, 66–79. 28. Eaton, F.M. Chapter: Chlorine. In Diagnostic Criteria for Plants and Soils; Chapman, H.D., Ed.; University of California: Riverside, 1966; 98–135. 29. Wang, D.Q.; Guo, B.C.; Dong, X.Y. Toxicity effects of chloride on crops. Chin. J. Soil. Sci. 1989, 30, 258–261. 30. Whitehead, D.C. Chlorine deficiency in red clover grown in solution culture. J. Plant Nutr. 1985, 8, 193–198. 31. Parker, M.B.; Gaines, T.P.; Gascho, G.J. The chloride toxicity problem in soybean in Georgia. In Special Bulletin on Chloride and Crop Production, 2nd Ed.; Jackson, T.L., Ed.; Potash & Phosphate Institute: Atlanta, 1986; 100–108. 32. Yang, J.; Blanchar, R.W. Differentiating chloride susceptibility in soybean cultivars. Agron. J. 1993, 85, 880–885. 33. Parker, M.B.; Gaines, T.P.; Gascho, G.J. Chloride Effects on Corn. Commun. Soil Sci. Plant Anal. 1985, 16, 1319–1333. 34. Yin, M.J.; Sun, J.J.; Liu, C.S. Contents and distribution of chloride and effects of irrigation water of different chloride levels on crops. Soil Fert. 1989, 1, 3–7 (Chinese). 35. Robinson, J.B. Fruits, vines and nuts. In Plant Analysis—An Interpretation Manual; Reuter, D.J., Robinson, J.B., Eds.; Inkata Press: Sydney, 1986; 120–147. 36. Corbett, E.G.; Gausman, H.W. The interaction of chloride and sulfate in the nutrition of potato plants. Agron. J. 1960, 52, 94–96. 37. James, D.W.; Weaver, W.H.; Reeder, R.L. Chloride uptake by potatoes and the effects of potassium chloride, nitrogen and phosphorus fertilization. Soil Sci. 1970, 109, 48–53. 38. Kafkafi, U.; Valoras, N.; Letay, J. Chloride interaction with NO3 and phosphate nutrition in tomato. J. Plant Nutr. 1982, 5, 1369–1385.

Classification Systems: Australian and New Zealand Terry A. Isbell A. E. Hewitt Landcare Research, Lincoln, New Zealand

INTRODUCTION Although Australia and New Zealand are relatively close neighbors, both are independent island states and this probably accounts for some degree of isolation between the two countries. In terms of size, climate, geology, vegetation, and land forms, there are striking differences, and it is not surprising, that the New Zealand soil pattern is mostly very different from that of much of Australia. It also follows for these factors, and other national differences, that separate soil classification schemes have always been in use. It is widely accepted that classification is an essential part of any scientific discipline and is a necessary part of the language of science. No classification scheme can remain static; as new knowledge is gained, soil classifications need to be updated and improved. The history of national soil classifications in Australia[1,2] is a good example of how benefits are obtained from modifications in the light of new knowledge. This, in effect, largely explains why Australia has rightly had a succession of national classification schemes. It is also of interest to note that two of the still commonly used schemes—the Factual Key of Northcote[3] and the so-called Handbook of Australian Soils[4]—were a direct result of the decision to hold the Ninth International Congress of Soil Science in Adelaide, Australia, in 1968. This meeting served as a catalyst to increase rapidly the knowledge about Australian soils by means of a targeted approach to the mapping of the continent at a published scale of 1 : 2 million using the Factual Key.[3] This was achieved on time in spite of the size and soil diversity of the Australian continent. The Factual Key is a bifurcating, hierarchical scheme with five categorical levels. All classes are mutually exclusive, and the keying morphological attributes are determined in the field, including the pH and the presence of carbonate as determined by a simple field test. Most class names are descriptive, e.g., ‘‘Hard Pedal Mottled Red Duplex Soils.’’ The associated Handbook of Australian Soils[4] was a separate exercise to the soil mapping project and is a compendium of morphological and laboratory data of 94 soil profiles which were visited on the 1968 Congress field excursions, plus some extra representative soils (147) not visited in the field tours. The ‘‘Handbook’’ 230 Copyright © 2006 by Taylor & Francis

is not a formal classification but is an assemblage of ‘‘great soil groups’’ largely derived from Stephens.[5] This was more or less current at the time and was essentially based on the earlier United States Department of Agriculture schemes that were in vogue prior to the formal advent of the early Soil Taxonomy Approximations of the 1960s. Both the Factual Key and the Handbook are still widely used in Australia, the former mainly because of the Australia-wide map coverage at 1 : 2 million, while the Handbook is still very useful because of the detailed accounts of the morphology, micromorphology, and laboratory data for a large number of widespread Australian soils. A disadvantage of the Handbook is the lack of a key to the soil groups. The most comprehensive of the earlier New Zealand national soil classification systems is the New Zealand Genetic Soil Classification[6] reviewed by Hewitt[7] in 1992. In his classification, Taylor[6] developed what would now be called soil landscape models, in which soil groups were related to environmental factors employing the zonal concepts of the Russians. The classification by Taylor was first published in 1948 as a legend to the first National Soil Map, but with time, a number of weaknesses became apparent. Examples include the failure of the genetic scheme when applied at the scale of the farm paddock and the difficulty of soil correlation between regions. It became obvious that a new model was required that made provision for important classes of New Zealand soils.

NATIONAL SOIL CLASSIFICATION SYSTEMS IN THE MODERN ERA In the 1970s and 1980s, there was an increase in soil surveys in some Australian states, particularly Queensland and the Northern Territory, where the then-current Northcote and Stace, et al. systems were often found to be inadequate to cater for many ‘‘new’’ Australian soils, particularly in the wet tropics and in the subtropics. A soil classification committee with an Australia-wide charter was set up in the early 1980s to remedy the lack of an adequate, up-to-date national soil classification. Extensive field travel was carried out in most parts of Australia to gather field Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042642 Copyright # 2006 by Taylor & Francis. All rights reserved.

Classification Systems: Australian and New Zealand

231

Fig. 1 Schematic summary of the 14 orders of the Australian soil classification. (From Ref.[10].) (Note: this figure is not to be used as a key.)

and laboratory data to build up a database, which would eventually contain some 14,000 published and unpublished profiles, many with laboratory data; mainly chemical. The ‘‘First Approximation’’ of a new national scheme was first compiled in 1989 and was widely circulated in Australia for comment. Two further approximations were produced and circulated before the published version appeared in 1996 titled The Australian Soil Classification (ASC).[2] It should be noted that due assessment was made of two so-called international systems viz., Soil Taxonomy[8] and the subsequent revisions of most of its Orders, and the other major ‘‘international’’ system that of the FAO–UNESCO (1990) Soil Map of the World and its updated (1998) version titled World Reference Base for Soil Resources.[9] While these two so-called international schemes are of considerable value and interest for comparative purposes, it is more likely that most Australian pedologists would agree on the need for an appropriate national system that is


regularly updated. The possibility of using methods of numerical classification was also examined, but was thought to be inappropriate because of the great variability and inconsistencies in the soil datasets, e.g., lack of representation of important groups of soils.

The Australian Soil Classification This scheme[2,10] is now widely used in most Australian soil surveys and pedological studies, primarily in an identification role. The scheme may be described as a multicategoric, hierarchical (order, suborder, great group, subgroup, and family), general purpose scheme with classes defined on the basis of diagnostic horizons or materials and their arrangement in vertical sequence as seen in an exposed profile. The classes are mutually exclusive and the allocation of new or unknown individuals to the classes is by means of keys. The scheme is

232

open-ended and new classes can be defined as knowledge increases, although these may not necessarily be added in sequence to the present list. Fourteen orders are currently defined (Fig. 1). A recent innovation is the development of an interactive key to The Australian Soil Classification available on CD-ROM.[11] The structure of the ASC has strong similarities with that of Soil Taxonomy, and although the emphasis is on field morphology, laboratory data have been used as appropriate. All technical terms are defined in a glossary. In a companion publication,[10] further descriptions of diagnostic horizons and materials as well as the general occurrence of each order are given except for Anthroposols where data were unavailable. The major classes of each order and the use of major attributes in the subdivision of the orders are discussed. The main chemical properties used are pH and exchangeable cations, with pH commonly being determined in the field. A table is provided giving approximate correlations between ASC orders, three other Australian classifications, and Soil Taxonomy orders. Perhaps the major benefit of the ASC is that it provides a means for efficient communication and a systematic framework for understanding and learning about the properties and inter-relationships of soils. Experience with the Factual Key[3] suggested retention of the use of color and generalized texture profiles at high levels in the hierarchy of the new system, although formal justification of such decisions is often difficult to demonstrate. Another feature of the ASC that partly overcomes some of the problems associated with hierarchies is the use of family criteria at any level of the system.


one (order, group, subgroup, and series), with class distinctions based on diagnostic horizons and materials. Keys are provided to enable easy class definitions and identification. Total number of orders is 15: Allophanic Soils, Anthropic Soils, Brown Soils, Gley Soils, Granular Soils, Melanic Soils, Organic Soils, Oxidic Soils, Pallic Soils, Podzols, Pumice Soils, Raw Soils, Recent Soils, Semiarid Soils, and Ultic Soils. Accessory properties of the orders as well as concept, occurrence, and correlation with the earlier New Zealand Genetic Soil Classification and appropriate classes of Soil Taxonomy are also listed. The scheme is open-ended and new classes can be incorporated into the hierarchy.

CONCLUSIONS This brief review of the development of two national soil classification systems indicates the need for some taxa which are required for one country but not the other. Of the 15 New Zealand soil orders that have been defined, at least two (Allophanic Soils and Pumice Soils) are virtually absent in Australian soil landscapes. Conversely, some widespread Australian soils are only partly represented in the New Zealand system. Particular examples are the widespread Australian semiarid soils characterized by profiles which feature a clear or abrupt change to a textural B horizon, which is frequently sodic. Finally, there are some soils common to both countries that probably could be satisfactorily classified at the order level by the other system. A few such examples include Arthropic soils and Arthroposols, Gley soils and Hydrosols, Organic soils and Organosols, Oxidic soils and Ferrosols, and Podzols and Podosols.

New Zealand Soil Classification With the advent of Soil Taxonomy, New Zealand soil scientists became heavily committed to its international development, involving a five-year testing program in New Zealand. As in Australia, difficulties became apparent, particularly with regard to its complexity. The results of the investigations showed that Soil Taxonomy made inadequate provision for important classes of New Zealand soils, particularly in the case of Inceptisols. Much of the new order of Andisols was based on New Zealand, where Guy Smith spent 12 months studying these and other soils. The inadequacies of Soil Taxonomy and the older New Zealand genetic classification to serve as an up-to-date national classification led to the development and publication of new material, largely through the efforts of Hewitt.[7,12–15] The most recent version[13] has a helpful introductory section in which concepts, objectives, and principles are outlined. The scheme is a hierarchical


REFERENCES 1. Isbell, R.F. A brief history of national soil classification in Australia since the 1920’s. Aust. J. Soil Res. 1992, 30, 825–842. 2. Isbell, R.F. The Australian Soil Classification; CSIRO Publishing, 1996. 3. Northcote, K.H. A Factual Key for the Recognition of Australian Soils, 4th Ed.; Rellim Technical Publications: South Australia, 1979. 4. Stace, H.C.T.; Hubble, G.D.; Brewer, R.; Northcote, K.H.; Sleeman, J.R.; Mulcahy, M.J.; Hallsworth, E.G. A Handbook of Australian Soils; Rellim Technical Publications: South Australia, 1968. 5. Stephens, C.G. A Manual of Australian Soils, 3rd Ed.; CSIRO: Melbourne, 1962. 6. Taylor, N.H. Soil Map of New Zealand, 1:2,027,520 Scale; DSIR: Wellington, 1948. 7. Hewitt, A.E. Soil classification in New Zealand: legacy and lessons. Aust. J. Soil Res. 1992, 30, 843–854.


8. Soil Survey Staff. In Soil Taxonomy, a Basic System of Soil Classification for Making and Interpreting Soil Surveys; USDA Agriculture Handbook No. 436; U.S. Government Printing Office: Washington, 1975. 9. FAO. World Reference Base for Soil Resources; FAO World Soil Resources Report 84, 1998. 10. Isbell, R.F.; McDonald, W.S.; Ashton, L.J. Concepts and Rationale of the Australian Soil Classification, ACLEP; CSIRO Land and Water: Canberra, 1997. 11. Jacquier, D.W.; McKenzie, N.J.; Brown, K.L.; Isbell, R.F.; Paine, T.A. The Australian Soil Classification: An Interactive Key; CSIRO Publishing, 2001.


233

12. Hewitt, A.E. New Zealand Soil Classification (Version 2.0); DSIR Division of Land and Soil sciences Technical Record DN 2; Department of Scientific and Industrial Research: Wellington, 1989. 13. Hewitt, A.E. New Zealand Soil Classification; Landcare Research Science Series No. 1; Manaaki Whenua Press: Lincoln, 1998. 14. Hewitt, A.E. Methods and Rationale of the New Zealand Soil Classification; Landcare Research Science Series No. 2; Manaaki Whenua Press: New Zealand, 1993. 15. Clayden, B.; Webb, T.H. Landcare Research Science Series No. 1; Manaaki Whenua Press: Lincoln, New Zealand, 1994.

Classification Systems: French Jean-Paul Legros Science du Sol, INRA, Montpellier, France

INTRODUCTION France is a very small country. But it was previously the head of an empire covering a part of Africa and Asia. During the International Exhibition of 1900, in Paris, Dokoutchaev explained his concepts concerning pedology, attracting the interest of the French scientists on this subject. The two facts explain the interest of French scientists in soil classification.

BRIEF HISTORY From the beginning of pedology, the French scientists used genetic classifications of soils in relation to the work of the Russian school of Dokuchaiev (1846– 1903) and Glinka (1867–1927). For example, Vale´rien Agafonoff (1863–1955), a Russian soil scientist, came to Paris fleeing his country during the revolution of 1917. He built one of the first soil maps of France with the corresponding legend (created in 1928, but published mainly in 1936 after improvements).[1–3] Using the previous and successive works of Lagatu (1862–1942),[4,5] Demolon (1881–1954), Kubieńa (1897–1970), Erhart (1898–1982) and several German authors, Georges Aubert (born in 1913), and Philippe Duchaufour (1912–2000) presented a first French classification in Paris, in 1956, during the Sixth International Congress of Soil Science. A second version of this system was presented in 1962 and is known as ‘‘Aubert-Duchaufour classification.’’ In 1960, Duchaufour published his famous ‘‘Prećis de Pe´dologie,[6]’’ which popularized his classification system to a much wider audience. Then, in 1967, the whole French scientific community collaborated and completed the system which became the ‘‘Classification of the CPCS’’ (Commission de Pe´dologie et de Cartographie des Sols). This document was used for 20 yrs and popularized a vocabulary presently used in the ordinary pedological language of French scientists, e.g., sol brun, ranker, rendzine, etc. During this period, the USDA system and the FAO Soil Legend were developed (1960 and 1981). Simultaneously, the development of Computer Science and Statistical Science demonstrated that the soils could be classified in a more objective way by searching for Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042644 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

the mathematical similarities between taxa and objects rather than trying to enter a soil in a more or less adapted taxon of a rigid classification system. In this context, since 1986, the French community met, worked, and presented a first version of ‘‘Re´fe´rentiel Pe´dologique’’ (RP) in 1992. A second version was produced in 1995,[7] with an English translation in 1998,[8] and followed by Italian and Russian translations (2000). Because several French scientists were involved in the development of both the new French trials and the World Reference Base (WRB),[9] the two systems have many similarities in both philosophy and organization, although not necessarily in the details of vocabulary. Thus, it is logical to review the RP in comparison with the WRB.

COMPARING THE RP AND THE WRB Principal points of comparison are the following:

The 40 diagnostic horizons of the WRB are replaced by the 73 ‘‘Horizons de Re´fe´rence’’ of the RP.

The 30 reference soil groups of the WRB are replaced by the 30 ‘‘Grands Groupes de Re´fe´rences’’ (major groups of references or ‘‘GER’’). But these 30 categories do not match exactly because, in the French system, the tropical soils are not considered. This last point is regretted[10] (10), considering the large vast experience of French scientists in Africa. Theses 30 Groupes are divided into 102 ‘‘Sols de Re´fe´rence.’’

The WRB ‘‘Diagnostic properties,’’ ‘‘Diagnostics materials,’’ ‘‘Formative elements,’’ and ‘‘Prefixes’’ correspond to the ‘‘Qualificatifs’’ (Qualifiers) of the RP. In the RP, all these elements are grouped in a single list of 235 terms. The more original point of the RP is that a Diagnostic horizon alone is in general, not sufficient to recognize a Re´fe´rence. Several Diagnostic horizons are associated to identify a Solum (profile development as A, E, B, C including a part of the underlying rock).[11] For this reason, the ‘‘Horizons de Re´fe´rence’’, in the RP, are defined considering strongly their relative development (A, B, C horizons, 247

248

Classification Systems: French

etc.). This indicates clearly that a part of the previous genetic approach remains inside the RP. Its organization is presented in Fig. 1. Even if it is not fully completed, the RP, richer in taxa and qualifiers for a smaller part of the world, may be considered to be more precise than the WRB. It is organized in three logical levels that are not truly hierarchical. It allows the user to identify a soil as an intermediate between two taxa, e.g., a ‘‘REGOSOL– CRYOSOL’’. Table 1 shows the approximate links between WRB and RP at the level of GER, Re´fe´rences and Groups. The most specific Re´fe´rences of the RP if we compare with WRB are presented in Table 2 with short explanations. All this demonstrates that it is very easy to go from one system to the other at the level of Groups and Re´fe´rences even if ‘‘Calcisol’’ is a false cognate. (Calcisol means calcareous in WRB but saturated and not calcareous in the RP.) The situation is more complicated at the level of soil types. Many of the qualifiers seem similar, but this may not truly be the case. Table 3 shows some of the difficulties which may be encountered. Table 4 characterizes the French soils using the RP system. The data were kindly provided by the Service d’Etude des Sols et de la Carte Pe´dologique de France (Orleáns).

Table 1 Relationships between the WRB groups (left) and the GER or the Re´fe´rences of the RP (right) Re´fe´rentiel pe´dologique

WRB Histosol

Histosol

Cryosol

Cryosol

Anthrosols

Anthroposols

Leptosols

Lithosols (superficial) Rankosols (on acid rocks) Organosols (rich in organic matter) Rendosols (on calcareous rocks)

Vertisols

Leptismectisols (A=C or A=R profile) Vertisols (with B horizon)

Fluvisols

Fluviosols Thalassosols (estuarine=marine soils)

Solonchaks

Salisols

Gley soils

Re´ductisols (dominant reduction) Re´doxisol (dominant oxidation)

Andosols

Andosols Vitrosols (with glass)

Podzols

Podzosols

Plinthosols

Not already studied

Ferralsols

Not already studied

Solonetz

Sodisols

Planosols

Planosols

Chernozems

Chernosols

Kastanozem

DISCUSSION To pass from a purely genetic classification to an international reference system, such as the WRB, was rather a long journey for the French pedological community.[12] The construction of the RP, initiated by Denis Baize and Michel-Claude Girard, was a good way to test the possibility of this large change in the French method of thinking.

Fig. 1 Organization of the Re´fe´rentiel Pe´dologique.


Phaeozem

Phaeosols

(Greyzems)

Grisols

Gypsisols

Gypsosols

Durisols

Not already studied

Calcisols

Rendosol (rendzine) Rendisols (same morphol., saturated) Calcosols (calcareous with B) Dolomitosols (with MgCO3) Calcisols (noncalcareous, saturated) Magnesisols (with Mg2þ on clay) Calcarisols (calcaric within 25 cm)

Albeluvisols

Luvisols (for a part)

Alisols

Not already studied

Nitosols

Fersialsols (for a part)

Acrisols

Not already studied

Luvisol

Luvisol

Lixisols

Not already studied

Umbrisols

Alocrisols humiques, rankosols

Cambisols

Brunisols Pelosols (rich in clay but not 2=1)

Arenosol

Arenosol

Regosol

Re´gosol

Classification Systems: French

249

Table 2 Specific Re´fe´rences in the RP Alocrisols

Table 4 Inventory of the French soils using the RP (from INRA-SESCPF)

Acidic but without argic horizon (i.e., different from the WRB ‘‘Acrisols’’)

Colluviosols

Re´fe´rentiel Pe´dologique

From colluvium, i.e., on slopes, in parallel with fluviosols From Peyre ¼ stone in some local French languages, i.e., with important coarse fraction

Peyrosols

Veracrisols

Sort of Acrisols rich in earth worms

Recently (year 2005, June), the French community of Soil Science fall in agreement (vote in an Administrative Council of AFES) on the following ideas:1) the RP system will be updated adding the tropical soils that were so precisely studied in the period of the French African Empire and 2) this new version of the RP will be built taking in consideration the WRB to get a full compatibility between the two systems. Working in such a way, French scientist will conserve their national system allowing them to focus on such taxa for regional reasons, but preserving the international dialog through the WRB.

Corresponding WRB group

% of France

Calcosols, Brunisols

48.5

Cambisols

Luvisols

14.5

Luvisols

Rendosols

8.3

Calcisols (p.p)

Fluviosols

7.8

Fluvisols

Podzolized luvisols

6.4

Albeluvisols

Podzosols

5.6

Podzols

Lithosols

2.3

Leptosols (p.p)

Rankosols

1.8

Leptosols (p.p)

Andosols, vitrosols

1.0

Andosols

Areńosols

0.8

Arenosols

Re´doxisols, Re´ductisols

0.4

Gleysols

Re´gosols

0.4

Regosols

Salisols

0.4

Solontchaks

Histosols

0.3

Histosols

Phaeosols

0.1

Phaeozems

Vertisols

0.1

Vertisols

Planosols

0.1

Planosols

Autres sols et surfaces

1.2

Others soils and surfaces

CONCLUSIONS in the RP, at an international level, seems limited because of its great similarity with the WRB.

For the French scientific community, the development of the RP is a good opportunity to work together on the concepts of the soil classifications and to study the genesis and the functioning of the soils identified in France. Moreover, in the RP volume, the texts that present the Andosols (the soils with hydromorphic features, and the different kinds of humus) are valuable contributions to Pedology. Nevertheless, the interest

ARTICLE OF FURTHER INTEREST World Reference Base for Soil Resources, p. 1918. Classification: Need for Systems, p. 227.

Table 3 Differences between qualifiers in WRB and RP Examples Case

Problem

WRB terminology

RP terminology

1

Term specific of one of the two systems (scarce case)

Carbic

Clinohumic (isohumic)

2

Same meaning but different terms (scarce case)

Alumic (Al sat >50%)

Aluminic (Al

3

Same name but slightly different meanings (general case)

Magnesic Ca2þ=Mg2þ 0.84 mm in diameter, resists erosion from all but the highest winds.[9] Clods of this size are not easily moved by wind, and they protect smaller clods and particles in their lee. Loams, silt loams, and clay loams tend to form the most stable aggregates and are, therefore, the least affected by erosive winds. The type of tillage equipment used has a definite influence on soil cloddiness and surface roughness. Smika and Greb[15] found that tillage by machines other than the chisel tend to reduce the non-erodible soil aggregation. One-way, offset or tandem disks leave a smooth surface. Subsurface sweeps, because they do not disturb the soil surface, do not create a rough, ridged soil surface, but they do maintain a greater vegetative roughness by allowing some of the vegetation to remain erect.[16]

ROUGHENING THE LAND SURFACE Soil surface roughness is composed of anchored vegetative material, soil ridges, soil clods, or combinations of all three. All help to control wind erosion by lowering the wind velocity near the soil surface and by sheltering erodible soil fractions.[17] Tillage implements form ridges and depressions which alter wind velocity. The depressions behind the ridges trap saltating soil particles and stop the normal build-up of eroding material downwind. Emergency Tillage Emergency tillage is a last-resort wind erosion control practice that can provide a rough, cloddy surface. It is usually carried out when vegetative cover is depleted by excessive grazing, drought, improper or excessive tillage, or by growing crops that produce little or no residue (Fig. 6). Emergency tillage is an inadequate wind erosion control measure and its only purpose is to create a temporary erosion-resistant soil surface. Implements such as listers, chisels, shovels and ‘‘sandfighters’’ should traverse fields at right angles to erosive winds to roughen the soil surface and bring clods to the surface.[9] Listers and narrow chisels were found by Chepil and Woodruff[16] to have the most effective tillage points for emergency tillage. Listers provide a high degree of roughness, and in extremely sandy soils,


601

Fig. 6 Field following emergency tillage. (Courtesy of USDA.)

where clods can be produced only by deep tillage, they are the most effective tools available. Chisel cultivators are more widely used because they require less power and destroy less growing crop than listers.

RESHAPING THE LAND TO REDUCE EROSION ON KNOLLS Reshaping the land by leveling knolls and benching slopes to shorten the unsheltered distance is an option in wind erosion control, but is usually not economical or practical. Because land reshaping is very costly, other effective control measures, such as no-till and seeding to permanent grass, are usually options that are more viable. Hills and knolls affect tillage system requirements indirectly by influencing wind shear stress. When the wind blows over a hill, streamlines of airflow are squeezed together, which increases the wind velocity and shear stress, thereby increasing the erosion potential on the windward slope and hilltop. Consequently, this increases the amount of residue, cloddiness, or roughness needed to control wind erosion on the knoll.

OTHER WIND EROSION SITUATIONS Wind Erosion on Irrigated Land Wind erosion on irrigated land can be a serious problem in areas characterized by variable high-wind velocities, where the soils are organic or quite sandy and low in organic matter or where crop residues are inadequate. Under certain conditions it is impractical and wasteful of water to irrigate frequently enough to prevent a finely pulverized surface soil from blowing.

602

The depth of drying may only be a fraction of an inch and the soil below this may be wet, but if the immediate surface is dry and the wind is strong enough, the top layer can erode unless the soil particles are consolidated into clods or protected by vegetation. The basic element in erosion control by tillage on irrigated land, as on dryland, is the creation of a rough, cloddy surface which will resist the force of the wind, decrease its velocity at the ground level and trap moving soil. Sandy soils, usually found in irrigated areas, are far more difficult to protect by emergency measures than fine-textured soils. Wind Erosion Control on Sand Dunes and Other Problem Areas Wind erosion control on sand dunes and other problem areas require measures that are more intensive to get sand dunes in check. Dunes lack a soil profile because they are unstable and underdeveloped. The sand is fine, loose and easily moved by wind. It has no organic matter, and consequently retains little moisture for plants and has inherently low fertility. Sand dunes and drift areas often require artificial barriers or cover for stabilization before vegetation can be established. These include oil, clay gravel, picket fence, brush, straw, and hay. Clay is effective, but is expensive. Hay or straw can be used as temporary mulches on blowout or small areas of dune sand at road cuts, around dwellings and other disturbed areas. They provide some organic matter, which is critical for successful dune plantings. The establishment of permanent vegetation is the final objective in the stabilization of dunes. Other Non-Vegetative Erosion Protection Some of the non-vegetative and processed vegetative materials used are gravel and crushed rock, various surface films such as resin-in-water emulsion (petroleum origin), rapid-curing cutback asphalt, asphaltin-water emulsion, starch compounds, latex-in-water emulsion (elastomeric polymer emulsion), by-products of the paper pulp industry, and wood cellulose fiber.[2] Several of these spray-on adhesives are available for temporary wind erosion control of vegetable seedlings on mineral soils. Some of the adhesives are relatively expensive, but a few are economically feasible on high-value crops threatened by serious blowing that cannot be controlled by other methods.


Erosion by Wind: Control Measures

REFERENCES 1. Lyles, L. Basic wind erosion processes. Agric. Ecosystems Environ. 1988, 22=23, 91–101. 2. Woodruff, N.P.; Lyles, L.; Siddoway, F.H.; Fryrear, D.W. How to Control Wind Erosion; U.S.D.A., SEA Agric. Inf. Bull. No. 354; U.S.D.A., A.R.S.: Washington, D.C, 1972; 22 pp. 3. Greb, B.W. Reducing Drought Effects on Cropland in the Western-Central Great Plains; U.S.D.A., SEA Agric. Inf. Bull. No. 420; Washington, D.C, 1979; 31 pp. 4. Haas, H.J.; Willis, W.O.; Bond, J.J. Summer Fallow in the Western United States; Agriculture Research Service Conservation Research Report No. 17; U.S.D.A.: Washington, D.C, 1974; 149–160. 5. Watson, S., Ed.; Kansas No-Till Handbook; Kansas State University: Manhattan, KS, 1999; 72 pp. 6. Rice, R.W. Fundamentals of No-Till Farming; American Association for Vocational Instructional Materials: Winterville, 1983; 148 pp. 7. Crovetto, C. Stubble over the soil. In The Vital Role of Plant Residue in Soil Management to Improve Soil Quality; Amer. Soc. Agronomy: Madison, WI, 1996; 245 pp. 8. Domitruk, D., Crabtree, B., Eds.; Zero Tillage. Advancing, the Art; Manitoba-North Dakota Zero Tillage Farmers Association: Brandon, Manitoba, 1997; 40 pp. 9. Tibke, G. Basic principles of wind erosion control. Agric. Ecosystems Environ. 1988, 22=33, 103–122. 10. Skidmore, E.L.; Hagen, J.L. Reducing wind erosion with barriers. Trans. ASAE 1977, 20, 911–915. 11. Siddoway, F.H.; Fenster, C.R. Soil conservation: western great plains. In Dryland Agriculture; Dregne, H.E., Willis, W.O., Eds.; Monograph No 23; Amer. Soc. Agronomy: Madison: WI, 1983; 231–246. 12. Chepil, W.S. Wind erosion control with shelterbelts in North China. Agron. J. 1949, 41, 127–129. 13. Black, A.L.; Bauer, A. Soil conservation: northern great plains. In Dryland Agriculture; Dregne, H.E., Willis, W.O., Eds.; Monograph No 23; Amer. Soc. Agronomy: Madison, WI, 1983; 247–257. 14. Aase, J.K.; Siddoway, F.H.; Black, A.L. Effectiveness of grass barriers for reducing wind erosiveness. J. Soil Water Conserv. 1985, 40, 354–360. 15. Smika, D.E.; Greb, B.W. Nonerodible aggregates and concentration of fats, waxes and oils in soils as related to wheat straw mulch. Soil Sci. Soc. Am. Proc. 1975, 39, 104–107. 16. Chepil, W.S.; Woodruff, N.P. The physics of wind erosion and its control. Adv. Agron. Water Conserv. 1963, 15, 211–302. 17. Armbrust, D.V.; Chepil, W.S.; Siddoway, F.H. Effects of ridges on erosion of soil by wind. Soil Sci. Soc. Am. Proc. 1964, 28, 557–560.

Erosion by Wind: Effects on Soil Quality and Productivity John Leys Department of Land and Water Conservation, Gunnedah, New South Wales, Australia

INTRODUCTION

Soil Texture Changes

Accelerated erosion on agricultural lands has adverse effects on soil quality and productivity through the removal of soil particles and nutrients. The majority of reviews of this topic have concentrated on the impact of water erosion on soil and crop productivity.[1,2] However, there is a growing number of research on the impact of wind erosion on soil quality and productivity, which is the focus of this entry. Wind erosion reduces soil quality and production using a mechanism different from that of water erosion. Water erosion removes the soil en masse, while wind erosion winnows the finer=lighter particles from the surface, leaving the larger (generally inert) particles behind. As a result, wind erosion removes topsoil[3] and reduces the soil clay and silt content[4] and the organic matter.[5] Wind erosion also has an impact on crop productivity. It sandblasts emerging crops,[6] reshapes the land surface, thereby making it difficult to traverse with wide agricultural implements, buries or undermines infrastructure such as fences and roads, and buries adjacent land with sand drift. This results in limiting the drifted land’s production in the short term.[7] Off-site impacts will not be discussed here, although wind erosion also has considerable off-site impact—e.g., reduces visibility,[8] deposits unwanted dust and off-farm associated contaminants off-farm, and raises airborne particulate levels,[9] with particle sizes less than 10 mm (PM10), which can have adverse health effects.[10]

The impact of erosion on soils has been measured over many years by using many different methods. Longterm analysis of soils exposed to wind erosion[11] showed a decline in the fertility and particle size distribution (PSD) of the surface soil. Over a 36-year period, a 6.5% increase in the sand fraction of the top 0–10 cm has been reported for midwestern U.S.A.[12] The comparison of the PSD of a soil that had been farmed for 30 years and had suffered periodic erosion with that of an adjacent soil under native vegetation,[13] revealed a loss of fines in the 10–100 mm fraction, as well as an increase in the coarse 350–1000 mm fraction (Fig. 1). Increases in the saltation fraction (>250 mm), and reduction in the 75–210 mm fraction, have been reported for the top 1 cm soil layer over a 15-week period for southeastern Australia.[3] Further research at the same site using a portable field wind tunnel indicated an increase in the dominant sediment population of PSD of the surface soil layer (approximately 0–500 mm depth) after a 30-min simulated erosion event. The analysis indicates that for the soil cultivation ridges, the proportion of the 300 mm sediment population increased from 60% in the parent soil to 86% after the simulated erosion. In the cultivation furrows, the 420 mm population increased from 0% to 85%. Analysis of the eroded sediments indicated that >78% of material being eroded fell within the 180 mm population, compared to 29% in the parent soil. The increase in the coarser fractions of the surface soil and the high proportion of finer fractions in the eroded sediment imply that wind erosion is winnowing the fines.

SOIL QUALITY Eroded Sediments Soil quality is generally adversely affected by wind erosion via the removal of soil fines (clay, silt, and organic fractions). Soil texture changes are largely irreversible, unless topsoil is imported to the site. Quantification of the changes in soil texture, as well as identifying the eroded fractions in the eroded sediments, indicates the magnitude of the decline in soil quality brought about by wind erosion. Descriptions of soil texture changes and the eroded sediments will be presented in the next two sections. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042680 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

A large number of studies in North America,[14] Belgium,[15] Nigeria,[16] Australia,[3] and China[17] show that the particle size decreases as the height increases, indicating a selective sorting of the eroded particles. When the vertical component of the wind exceeds the fall velocity of the eroded particles, they are removed from the site along with their associated nutrients. Australian results indicate that for a one-week period of monitoring, 27% of the total eroded sediment is in 603

604

Fig. 1 Particle size distribution of two adjacent soils. One has been farmed for 30 years and has undergone repeated erosion; the other is under native vegetation. The PSDs show an increase in the coarse fraction (arrow a) and a decrease in the finer fraction (arrow b) in the eroded soil.

the suspension fraction and thus, are removed from the eroded field.[3] These studies further highlight the winnowing action of the wind erosion processes and the potential loss of soil and nutrients.

SOIL PRODUCTIVITY When soil is eroded from a site, there is often a subsequent decrease in rooting depth and available water holding capacity.[13] There is also a loss of nutrients and a subsequent decline in soil productivity. So if there is wind erosion, where does the nutrient go and how much impact does it have on crop production? Soil Nutrient Loss There is considerable evidence that wind-eroded sediments are enriched with nutrients. The nutrient loss has been measured at the site of dust emission, at various heights above the eroded surface, and downwind sites of the eroded area, both immediately adjacent to drift banks of soil and from deposited dust. It is important to examine the magnitude of this nutrient loss. The nutrient content and particle size of eroded sediments collected at 0–0.5 m height during wind tunnel tests were similar to source sediments.[18] However, if the sediments are sieved and the 11.5% protein, nitrogen supply at the higher level of sufficiency and luxury range is economically preferred. Yet high soil nitrate-N fertility may lead


to excess of nitrate-N in soil and, consequently, to denitrification (greenhouse effect) and leaching. Excess nitrate-N in surface water bodies causes algal blooms and in drinking water causes methemoglobinemia in infants. Also, adverse environmental and health impacts occur when soil nutrient fertility is in toxic range (Fig. 2), especially for heavy metals, either because of human consumption of product or on ecosystems.

PERSPECTIVES Soil fertility evaluation has evolved in the last two centuries, from a major concern for enhancing crop production to that of maintaining environmental integrity. With increasing population, crop production will remain the main objective of managing the soil fertility. However, with the recent advances in multielement extractants and analysis and remote sensing and precision technologies for site-specific soil fertility evaluation, it is possible to meet the twin goals of economic optimum yields and minimal environmental pollution. By managing soil fertility this way, sustainable use of land, water, and nutrients is also ensured, resulting in enhanced soil and water quality for the current and future generations. In spite of the advances in sensor and analytical technologies, however, the challenge lies ahead in developing soil nutrient tests that closely mimic nutrient uptake by a crop.

CONCLUSIONS Significant progress has been made in soil fertility evaluation systems in last two decades because of valuable

Fertility Evaluation Systems

contributions made by ecologists and geographers, besides the continuing interest from agronomists and soil scientists. Consequently, a broad spectrum of technologies is being applied to soil fertility evaluation for sustainable natural resource management, food and fiber production, and ecologically sustainable practices for the benefit of both rural and urban communities. Further refinements in soil fertility evaluation systems would come from the integrative use of both classical and spectral technologies, and spatial and computational capacities for their appropriate applications to ecosystems and landscapes.

REFERENCES 1. Russell, E.W. Soil Conditions and Plant Growth, 10th Ed.; Longman: London, 1973; 49–50. 2. Mitscherlich, E.A. Das Gesetz des Minimums und das Gesetz des Abnehmenden Bodenertrages. Landwirtschaftliche Jahrbu¨cher 1909, 38, 537–552.


673

3. Robinson, D.J.; Reuter, J.B., Eds. Plant Analysis: An Interpretation Manual, 2nd Ed.; CSIRO Publishing: Melbourne, 1997; 1–572. 4. Black, C.A. Soil Fertility Evaluation and Control; Lewis Publishers: Boca Raton, FL, 1993; 1–746. 5. Havlin, J.L.; Beaton, J.D.; Tisdale, S.L.; Nelson, W.L. Soil Fertility and Fertilizers, 6th Ed.; Prentice Hall: Upper Saddle River, New Jersey, 1999; 1–499. 6. Peverill, K.I.; Sparrow, L.A.; Reuter, D.J., Eds. Soil Analysis: An Interpretation Manual; CSIRO Publishing: Melbourne, 1999; 1–369. 7. Dalal, R.C.; Hallsworth, E.G. Evaluation of the parameters of soil phosphorus availability factors in predicting yield response and phosphorus uptake. Soil Sci. Soc. Am. J. 1976, 40, 541–546. 8. Sanchez, C.A.; Couto, W.; Buol, S.W. The fertility capability soil classification system: interpretation, application and modification. Geoderma 1982, 27, 282–309. 9. Sudduth, K.A.; Hummel, J.W.; Birrel, S.J. Sensors for site-specific management. In The State of SiteSpecific Management for Agriculture; Pierce, F.J., Sadler, E.J., Eds.; ASA=CSSA=SSSA: Madison, 1997; 183–210.

Fertility: Environmentally Compatible Management Bal Ram Singh Agricultural University of Norway, Aas, Norway

INTRODUCTION Soil Fertility in the Past In writings dating back to 2500 B.C. the fertility of land is mentioned. Herodotus, the Greek historian, traveling through Mesopotamia some 2000 yr later reported the phenomenal yields obtained by the inhabitants of this land. Later Theophrastus (372–287 B.C.) recommended the abundant manuring of thin soils but suggested that rich soils be manured sparingly. During the seventeenth and eighteenth centuries, agricultural writings reflected that plants consisted of one substance, and most of the workers were searching for this principle of vegetation during this period. It was not until the later half of nineteenth and the beginning of the twentieth century that some progress was made to understand the subject of plant nutrition and crop fertilization. It was Justus von Liebig (1803–1873), a German chemist, who effectively deposed the humus myth and eventually developed the law of minimum. The law says that the growth of plants is limited by the plant–nutrient element present in the smallest quantity, all others being present in adequate amounts. These developments led to a rapid increase in chemical fertilizer’s use. Soil Fertility in Modern Times With advances in our knowledge with regards to various processes affecting the nutrient dynamics in soils, the definition of soil fertility is also refined. Soil fertility integrates the basic principles of soil biology, chemistry and physics to develop the practices needed to manage nutrients in a profitable and environmentally sound manner. The main focus of soil fertility is to manage nutrient status in soils to create optimum conditions for plant growth. Two fundamental principles underlay the study of soil fertility. First is the recognition that optimum nutrient status alone will not ensure soil productivity. Other factors, such as soil moisture and temperature, soil physical conditions, soil acidity and salinity, and biotic stress can reduce the productivity of even more fertile soils. Second is the realization that modern soil fertility practice must stress soil productivity and environmental protection.[1] Taking the second realization in perspective, it is imperative that soil 674 Copyright © 2006 by Taylor & Francis

fertility in relation to agricultural sustainability and environmental protection will be the main focus of this paper.

SOIL FERTILITY AND AGRICULTURAL SUSTAINABILITY In recent years, a new dimension to soil fertility in relation to agricultural production has been added. It is the concept of ‘‘sustainability,’’ which has been defined and interpreted differently by different workers. Okigbo[2] after analyzing the various definitions of sustainable agriculture by different workers defined ‘‘a sustainable agricultural production system as one that maintains an acceptable and increasing level of productivity that satisfies prevailing needs and is continuously adapted to meet the future needs for increasing the carrying capacity of the resource base and other worthwhile human needs.’’[2] The sustainability of a production system is location specific and is determined to a greater degree by an interaction among several production factors, viz. physiochemical (soil, climate, radiation etc.) biological (crop species, weeds and pests etc.), management and socioeconomic elements. Maintenance and management of soil fertility is central to the development of sustainable food production systems. Sustainability is dependent to a large degree on recycling the inputs into a production system, thereby increasing efficiency of output per unit of resource input. Soil fertility management is concerned with the essential plant nutrients, their amounts, availability to crop plants, chemical reactions in soil, loss mechanisms, processes making them unavailable or less available to crop plants, and ways and means of replenishing them in these soils.[3]

ESSENTIAL NUTRIENTS Because soil fertility involves management of nutrients required for plant production, it is important to describe briefly the elements, which are considered essential for plant growth. The essential nutrients required by higher plant are exclusively of inorganic nature, and Arnon and Stout[4] proposed the term essential nutrient (element). The essential element must meet three criteria to be Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001622 Copyright # 2006 by Taylor & Francis. All rights reserved.

Fertility: Environmentally Compatible Management

considered essential: 1) a given plant must be unable to complete its life cycle without the presence of mineral element in question; 2) the function of the element cannot be replaced by another element; and 3) the element is directly involved in the nutrition of the plant—for example as a component of an essential plant constituent such as enzyme—or must be required for a distinct metabolic step such as an enzyme reaction. Some elements, which either compensate for the toxic effect of other elements or they simply replace mineral elements in some of their less specific functions, such as maintenance of osmotic pressure, can be described as beneficial elements. Out of a large number of elements found in plants, 14 mineral elements are recognized as essential, whereas the requirement of chlorine and nickel is yet restricted to a limited number of plant species. The plant nutrients may be divided into macronutrients (N, P, S, K, Mg, Ca), micronutrients (Fe, Mn, Zn, Cu, B, Mo, Cl, Ni) and beneficial elements (Na, Si, Co). The last three elements have been found essential only for some plant species, for example, Na for plants with C4 photosynthetic pathway, Si for rice, and Co for fixation of atmospheric N by rhizobia and blue green algae, but they have not yet been included in the list of essential nutrients.[3]

675

Soil Quality The export of agricultural produce takes with it large amounts of nutrients (especially N, P, K, and S) and if these nutrients are not replenished, soil organic matter (SOM) and the fertility of the soil will decline and the soil will degrade. A number of such examples from the subsistence farming of the tropics are available, where ‘‘nutrient mining’’ of soils has been responsible for decline in soil fertility and soil quality leading to loss of productivity.[5] The rapid loss of SOM of the topsoil after clearing of the natural vegetation or under continuous cultivation is a common phenomenon in Africa. In Zambia it was found that the SOM content dropped from 59 to 32 Mg ha1 after 17 yr of cultivation of a newly cleared land, which represented an average loss of 1.6 Mg ha1 y1.[6] On the other hand, long-term studies carried out in temperate regions have shown that organic matter of soil can be maintained or raised modestly by proper fertilization and especially by organic manures and crop residue management.[7] Any management practice that results in an increase in the crop residues returned to soil will have a positive influence on soil organic matter,[7] which in turn will affect soil aggregation and related properties such as soil structure, erodibility, workability, and water infiltration. Growing of legumes is an effective way for promoting good soil aggregation and for reversing degradative trends.

MANAGEMENT OF SOIL FERTILITY AND THE ENVIRONMENT Water Quality As pointed out by Sims[1] that modern soil fertility in addition to involving soil productivity, must also include protection of the environment and thus the environmental aspects related to soil fertility are described under this section. Long-term use of organic and inorganic fertilizers to agricultural soils has led to slow build up of nutrient reserves and especially under temperate conditions. The same nutrients, which are considered essential for plant growth and crop production, if lost from the system, can create a concern for the environment. Increased use of fertilizers for meeting the world food demand and recycling of on farms (manures and slurries) and off farms organic wastes (municipal and industrial sludges) on agricultural lands in the last three decades have resulted in some undesirable effects on the environment in intensively cultivated areas. This has created greater awareness for environmental issues not only among scientists and policy makers but also among general public. Although there are a number of issues related to soil fertility and the environment, emphasis is placed on four main concerns of environmental protection, viz. soil quality, water quality, air quality, and heavy metals and food concerns.


The major concerns with regards to water quality are accelerated eutrophication of surface waters and nitrate (NO3) content of drinking water. Eutrophication, the rapid growth and decay of aquatic vegetation, is most often limited by P and sometimes by N concentration in water. The NO3 concentration of 10 mg L1 in drinking water is considered safe but higher concentration can cause methaemoglobinaemaia (reduced carrying capacity of blood for O2) in children. Both over fertilization and under fertilization can lead to N losses. Losses occur either as runoff to surface water or as leaching to underground waters. Under normal conditions, runoff losses in watersheds are low (e.g., 10 mg L1). Much of the N transported in runoff is particulate N associated with the sediments and it can range from 0 to 7 kg N ha1.[8] These losses can be reduced by incorporation of fertilizer and providing a good vegetation cover. Much of N losses occur through leaching but there is a large variation in the quantities of N lost in leaching depending on the amount of N fertilizer applied, soil type, crop grown and climatic conditions. Letey et al.[9] reviewed NO3 concentration in tile effluents from 55 sites of

676

California; the concentration ranged from 1 to 196 mg L1, with only one-quarter of the sites averaging 10 mg L1) nitrate levels.[10] From surface water viewpoint, P is the element of primary concern, since it is considered limiting for eutrophication. Excessive application of organic and inorganic fertilizers can result in building of P near the soil surface, which is subjected to soil erosion leading to P losses to surface waters. In aquatic ecosystems of southwestern Australia an excess of nutrients has caused serious eutrophication, which was manifested by excessive growth and accumulation of green and blue green algae.[11] Phosphorus is generally the limiting nutrient for algae growth and phosphatic fertilizers applied to nutrient-deficient, leaching, sandy soils were the main source of P in these ecosystems of Australia.

Air Quality The primary goal of good nutrient management and especially N should be to maximize N uptake at critical growth stages and to minimize transformation processes, which lead to formation of ‘‘greenhouse gases.’’ The main N-containing greenhouse gas is N2O. The N2O produced during nitrification and denitrification contributes to global warming and stratospheric ozone depletion. Annual losses of N as N2O from fertilized field soil can be as high as 40 kg h1 compared to that from unfertilized soils being 6.0 can help limit their release in the soil and availability to plants. A general advantage of organic fertilizers can be cost of the nutrient, but that is not always the case. It largely depends on the proximity of the organic fertilizer to the site of application and the value of the organic. In cases where a material must be placed in a landfill, the material may have a negative value due to the cost and maintenance of that landfill. If so, a material may be obtained at little or no cost, or even at a negative cost. As organic fertilizers generally have lesser nutrient values than modern manufactured inorganic fertilizers, the cost of transportation per unit of nutrient is often the controlling factor in value of the fertilizer. A second potential advantage of organic fertilizers is that they may provide nutrient release over an extended period of time; in particular, nitrogen (N) can be mineralized over a season to eliminate the need for repeated applications. Soil organic matter can be increased when organic fertilizers are applied at sufficient rates. The value of soil organic matter is well recognized, but the ability of added organic matter to substantially add to the pool of soil organic matter is debatable. It greatly depends on many factors, especially climatic conditions. General disadvantages of organics are their fertilizer ratios, that are not likely to be matched to plant needs, their nonuniform composition that can prohibit precise uniform applications, and their unpredictable release of nutrients, that may not coincide with the time a particular nutrient is most required by a particular plant. Some organic fertilizers may also contain unwanted, harmful, and=or toxic elements or compounds making them unsatisfactory for utilization. Some may be objectionable due to strong smells or irritants to humans. This latter disadvantage is now important to the quality of life in residential developments near sites of application.

ANIMAL RESIDUES Manures Animal manures were the first fertilizers used in agriculture. Feces and urine from large animals and feces Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001607 Copyright # 2006 by Taylor & Francis. All rights reserved.

Fertilizers: Organic

699

from poultry are the main sources. Most are applied, together with the bedding material, directly to the soil and usually in near proximity to the confined animals. Some is composted and=or pelleted to provide a material that is more agreeable for handling and transport. Nutrient composition can vary widely depending on the animal, its feed, the bedding, and the proportion of manure to bedding, and handling or storage prior to application (Table 1). Because of the importance of animal agriculture, manures will continue to be important sources of fertilizers and pose problems for agriculture and society. As farmers move toward precision applications (that provide improved nutrition without over-application that can result in pollution of soil or water), it becomes increasingly important to analyze manures and apply them in accordance with plant needs. Analyses provided indicate that phosphorus (P) composition is high relative to N for most plant needs. Therefore, application according to N needs provides too much P, that may be bioavailable and add to eutrophication of water via runoff. In some areas with high soil P, the recommendation is to apply manures on the basis of P needs. Such a philosophy will greatly limit applications in fields that are close to confined animals. It will also require N application and possibly other fertilizer nutrients from commercial fertilizers for balanced fertilization of most crops. One attempt to provide more uniform applications and products that are more environmentally friendly to transport is to compost or pelletize manures. Such operations concentrate the manure and reduce transportation costs. Many composting processes have been developed. To date the economics are favorable for some specialty plants but not for broad-scale crop production. Other Animal Residues Bone meal, dried blood, and other animal processing wastes are important fertilizers, but now account for little of the total fertilizer used in the USA. Bone meal is a good source of P, often used by the homeowner but seldom used in commercial agriculture. Dried blood is a rapid releasing organic N source. Other animal

wastes are variable in composition and are often considered as disposal problems rather than desirable fertilizers. Applications may be difficult due to poor and nonuniform physical characteristics. Chemical analysis should be known before application.

PLANT RESIDUES Leaving plant residues on the soil surface provides both nutrients for a following crop and soil protection from wind and rain erosion. Plant residue fertilization may also be obtained from the application of wastes or byproducts from many agricultural processing plants. Trash from cotton gins, filter mud from sugarcane, and processing wastes from fruits and vegetables are but a few of the plant residues utilized. Soil conservationists have long recommended keeping the soil protected by maintaining cover crops during the seasons when the main crops are not planted and returning those plant residues to the soil. The modern emphasis on conservation tillage includes planting directly into cover crops. If the cover crop is a legume, N may be supplied to the planted crop due to the N2-fixation from the legume. A good stand of alfalfa, clover, birdsfoot trefoil or vetch plowed into the soil may replace a fertilizer N application of 80–125 kg N/ha for a following crop of corn.[3,4] Presently, it is believed that such cover crops used in conservation tillage and not incorporated would supply nearly as much N to the following crop.

MUNICIPAL WASTES Sewage treatment plants produce fertilizers in the forms of effluent (the liquid portion) and as biosolids that are settled from the effluent. In some cases, biosolids are placed in a land fill, but costs of maintaining and monitoring see page from landfills is now making the economics favor transportation and land application. Proper handling in the treatment plant removes most human and animal health concerns, but not the stigma of such applications.

Table 1 Concentrations of N, P, and K on a dry-weight basis in commonly applied wastes Waste

N

P

K

References

Livestock manures

1–3%

0.4–2.0%

1–2.5%

[1,2]

Poultry manures

3–5%

1–3%

1–2%

[1,2]

Plant residues

1–7%

0.1–1.7%

0.1–9%

[1]

Municipal biosolids

2–9%

1.5–5%

0.2–0.8%

[9]

Municipal effluents

1.6–2.7 mg=L

0.2–1.2 mg=L

1.1–1.7 mg=L


[7,8]

700

Effluents Ammonium and nitrate-N should be monitored to avoid over-application of effluents, as N leaching can be a problem in sandy soils. When application is made year around, it is necessary to keep the soil covered with an actively growing crop to minimize leaching.[5] Phosphorus fixation capacity of the soil is an important determinant of the amount of effluent that should be applied without promoting a soluble-P problem. High fixing soils will handle greater amounts of P than soils with little fixing capacity.[6] Heavy metals are not a problem as they are concentrated in the solids. Nitrogen, P, and K concentrations of effluents vary quite widely (Table 1).

Biosolids Because of advances in technology, treatment plants are now able to remove nearly all heavy metals and bacteria, allowing the nutrient-rich organic material byproducts or biosolids to be recycled and applied as fertilizer. Elemental concentrations in biosolids from different sources have extremely wide variances (Table 1). They are generally good sources of both N and P for crops, but as for most manures, P application will be too great if enough is applied to satisfy N requirements. Application according to P requirements will allow little or no use in many cases.

OTHER BYPRODUCTS AND WASTES Many other organic byproducts and wastes may be suitable and even valuable as fertilizers. Included would be paper and pulp, vegetable and fruit process byproducts, food wastes, etc.—the list is too extensive to be included here. All wastes should be analyzed for


Fertilizers: Organic

nutrient elements and for potential pollutants prior to application.

REFERENCES 1. Huntley, E.E.; Barker, A.V.; Stratton, M.L. Composition and uses of organic fertilizers. In Agricultural Uses of Byproducts and Wastes; Rechcigl, J.E., MacKinnon, H.C., Eds.; American Chemical Society: Washington, DC, 1997; 120–139. 2. Miller, D.M.; Miller, W.P. Land application of wastes. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, FL, 2000; G-217–G-242. 3. Bruulsema, T.W.; Christie, B.R. Nitrogen contribution to succeeding corn from alfalfa and red clover. Agron. J. 1987, 79, 96–100. 4. Fox, R.H.; Piekielek, W.P. Fertilizer nitrogen equivalence of alfalfa, birdsfoot trefoil, and red clover for succeeding corn crops. J. Prod. Agric. 1988, 1, 313–317. 5. Hook, J.E. Comparison of the crop management strategies developed from studies at Pennsylvania state university. University of Minnesota, and the Muskegon county land treatment system. In Land Treatment of Municipal Wastewater; D’Itri, F., Ed.; Ann Arbor Science: Ann Arbor, MI, 1982; 65–78. 6. Sumner, M.E. Beneficial use of effluents, wastes, and biosolids. Commun. Soil Sci. Plant Anal. 2000, 31 (11–14), 1701–1715. 7. Ellis, B.G.; Erickson, A.E.; Jacobs, L.W.; Knezek, B.D. Crop management studies at the Muskegon county Michigan land treatment system. In Land Treatment of Municipal Wastewater; D’Itri, F., Ed.; Ann. Arbor Science: Ann Arbor, MI, 1982; 49–63. 8. Dowdy, R.H.; Clapp, C.E.; Marten, G.C.; Linden, D.R.; Larsen, W.E. Wastewater crop management studies in Minnesota. In Land Treatment of Municipal Wastewater; D’Itri, F., Ed.; Ann Arbor Science: Ann. Arbor, MI, 1982; 35–47. 9. Forste, J.B. Biosolids processing, products and uses. In Agricultural Uses of Byproducts and Wastes; Rechcigl, J.E., MacKinnon, H.C., Eds.; American Chemical Society: Washington, DC, 1997; 50–62.

Fertilizers: Urban Waste and Sludges Rory O. Maguire North Carolina State University, Raleigh, North Carolina, U.S.A.

James Thomas Sims University of Delaware, Newark, Delaware, U.S.A.

INTRODUCTION Environmental quality is a major issue in many parts of the world and includes important topics such as sustainability and environmental protection. Sustainability and environmental protection often overlap and cover issues such as soil conservation, careful use of finite resources such as inorganic fertilizers, and the responsible recycling of urban wastes and animal manures by agriculture. The beneficial reuse of urban wastes and by-products, e.g., sewage sludge (biosolids), has become an increasingly important issue to attain sustainability as the world’s population increases. This entry describes the uses of urban wastes and biosolids as fertilizers primarily for agricultural settings and the major trends, issues, options, and regulations involved. The main focus is on biosolids, as they constitute the largest portion of urban wastes that are currently land applied.

PRODUCTION OF URBAN WASTES USED AS FERTILIZER Biosolids Production and Quality There are several treatment processes that are routinely carried out in wastewater treatment plants. These determine the quality and properties of the biosolids produced, which in turn affects the options for land application. Following initial screening, wastewater can undergo primary, secondary, and tertiary treatments (Table 1). The treatment used at any particular wastewater treatment plant depends on the effluent discharge limits [e.g., biological oxygen demand (BOD) or nutrient content of the treated wastewater to be discharged into surface waters] and the proposed use for the biosolids and resources available. In addition to these wastewater treatment processes, de-watering of biosolids is frequently used to reduce biosolids volume, which can in turn reduce transportation and disposal costs.

TYPES OF URBAN WASTES THAT ARE USED AS FERTILIZERS

Composting of Biosolids and Municipal Solid Waste

Legislation in most countries requires treatment of wastewater from combined residential, commercial, and industrial sources. Treatment of wastewater produces a semisolid by-product that is commonly known as ‘‘sewage sludge’’ or ‘‘biosolids.’’ There are several disposal pathways for biosolids including incineration, landfilling, and application to land as a fertilizer. Use of biosolids in agriculture is a well-established and regulated process in many parts of the world, including U.S.A. and Europe. To a lesser extent, other urban and industrial by-products, e.g., paper waste, and municipal solid waste (MSW), such as leaves, grass clippings, or other organic materials, is also applied to agricultural land, often following composting. Sometimes biosolids can be cocomposted with industrial by-products or MSW.

Composting (often called cocomposting when two or more materials are composted together) is the decomposition of organic matter by micro-organisms in a controlled environment that has optimum moisture and oxygen contents. The increase in temperature during composting can destroy most pathogens. Composting of biosolids or the organic fraction of MSW may cause offensive odors and odor control systems, such as scrubbers and biofilters, are usually required in populated areas. Composting of biosolids involves mixing de-watered biosolids with a bulking agent, such as MSW, wood chips, or straw, followed by aerobic decomposition. Only a small proportion of biosolids and MSW are composted worldwide, but in certain areas composting accounts for a large proportion of the urban wastes generated. For example, Edmonton, Canada, is able to


705

706

Fertilizers: Urban Waste and Sludges

Table 1 Wastewater treatment processes Treatment name

Treatment process

Screening and grit removal

Generally a combination of sedimentation and passing through a large screen to remove stones, grit, and any large debris, such as branches, that may have entered the sewage system. The product of this process is nearly always regarded as a solid waste rather than biosolids and is usually landfilled

Primary

Follows screening and grit removal and usually involves sedimentation to remove suspended solids prior to secondary treatment

Secondary

Normally a biological treatment in which micro-organisms are used to reduce suspended solids and BOD, thereby eliminating fish kills when the wastewater is discharged. This is the minimum wastewater treatment process required in the U.S.A.

Tertiary

Used where higher standards are required for effluent quality or to improve biosolids quality for land application. Examples include addition of lime for pathogen and odor control, iron or aluminium salts for precipitation of P, and polymers for removal of suspended sediments

divert 70% of its residential waste from landfill using recycling and cocomposting.[1] Where the organic fraction of MSW is composted, it can either be source separated by the resident or screened and separated from regular residential garbage, and composted alone or mixed with biosolids before being composted.

REGULATIONS COVERING LAND APPLICATION OF URBAN WASTES Land application of biosolids normally depends on the wastewater treatment process used to produce the biosolids. As specific rules vary between countries, it is impossible to describe about all of them in this entry.[2,3] However, it is possible to make some generalizations about land application of biosolids. For example, secondary and=or tertiary treatment, mainly to control pathogens and odors, is normally required before land application is permitted. In U.S.A., biosolids applications are governed by Title 40 of the Code of Federal Regulations, Part 503,

commonly referred to as the ‘‘503 rule,’’ which sets limits for toxic metals in biosolids, and for both annual and cumulative loadings to land (Table 2). Limits for chromium and molybdenum are currently under consideration. The 503 rule also classifies biosolids as either ‘‘Class A’’ or ‘‘Class B’’ for pathogen control. The content of polychlorinated biphenyls in biosolids to be land applied is limited to a maximum of 50 mg=kg, under Title 40 of the Code of Federal Regulations, Part 761. Similar rules are in effect in the European Community, under Council Directive 86=278=EEC, implemented in 1986 and currently under revision.

AGRICULTURAL MANAGEMENT OF URBAN WASTES AS FERTILIZERS The benefits of biosolids and compost applications to soil quality are many and well documented. Biosolids and composts are good sources of nitrogen (N), phosphorus (P), and potassium (K). Typical biosolids

Table 2 Pollutant limits set by the 503 rule for toxic metals in biosolids applications to land

Pollutant Arsenic Cadmium

Ceiling concentration limits for all biosolids applied to landa (mg/kg)

Cumulative pollutant loading rate limits (kg/ha)

75

Annual pollutant loading rate limits (kg/ha/yr)

41

2

85

39

1.9

4300

1500

75

840

300

15

57

17

0.85

Nickel

420

420

21

Selenium

100

100

5

7500

2800

140

Copper Lead Mercury

Zinc a

Maximum concentration of pollutant permitted in any biosolids to be land applied.


Fertilizers: Urban Waste and Sludges

contain 4.0%, 2.0%, and 0.4% of N, P, and K, respectively, on a dry weight basis, while the equivalent values for MSWs are 0.7%, 0.2,% and 0.3%, respectively. Composted biosolids contain less total N than uncomposted biosolids, owing to the addition of other materials to aid composting and loss of ammonia during the composting process. The N in composted biosolids is released more slowly, which decreases nitrate leaching. Thus, the N is available to plants over a longer period, which is more consistent with crop N uptake patterns.[2] Biosolids and composts can also be good sources of Ca, Mg, and S, and of the micronutrients Fe, Cu, B, Zn, Mn, and Mo. Biosolids and composts can promote beneficial microbial activity and diversity that suppress plant diseases and the need for costly pesticides. Biosolids and composts can also improve water infiltration, water retention, and soil structure, which in turn increases resistance to wind and water erosion. In U.S.A., in 1998, land application of biosolids accounted for 41% of the total produced, while advanced treatments such as composting accounted for 12%.[2] The annual production of biosolids (7 million Mg=yr) is small in comparison to animal manure production (174 million Mg=yr) in U.S.A. However, the land application of biosolids constitutes a significant economic saving for many biosolids producers and saves landfill space, which is an increasingly expensive, finite resource in many areas.

POTENTIAL PROBLEMS ASSOCIATED WITH LAND APPLICATION OF BIOSOLIDS Legitimate concerns about the need to prevent toxic metals and pathogens from affecting human and ecosystem health have been addressed through regulation and record keeping of applications of biosolids to land. However, there is currently a debate as to whether the limits set for toxic metals and pathogens are strict enough. Some scientists think that metal bioavailability will increase as the organic matter added with the biosolids is mineralized, while others argue that the evidence to support this hypothesis is inconclusive. Biosolids and composts have a low N : P ratio compared to crop requirements. Biosolids applications are generally carried out according to N-based nutrient management plans that over-apply P, and can lead to a buildup of P in agricultural soils.[4] Buildup of soil P in many areas has been linked to increase in losses of P from agriculture to surface waters, with a corresponding decrease in water quality. In the future, biosolids may have to be applied according to P-based nutrient


707

management plans. This will decrease the amount of biosolids that can be applied per unit area of land, introduce the necessity for inorganic N fertilizers, and increase other associated costs. Nuisance issues such as odors and attraction of pests are of public concern when application sites are close to residential areas.[5]

CONCLUSIONS Efforts to recycle municipal wastes will likely continue, driven by human population growth, the continuing rise in cost of landfill, and the goal of sustainability in agriculture and the environment as a whole. However, the future of land-application programs for biosolids is uncertain. For example, the USEPA has forecast an increase in the beneficial use of biosolids in land-application programs because of increasing costs associated with landfill, and an increase in biosolids quality owing to stricter regulations covering biosolids production. Despite increasing biosolids quality, the argument over the long-term environmental impact of pollutants in biosolids will likely continue. Further regulation of the application of P to agricultural land may also increase costs associated with land-application programs for biosolids and municipal wastes. However, if the potential negative side effects associated with the land application of biosolids and composted waste are properly managed, then the beneficial use of these by-products will continue.

REFERENCES 1. www.gov.edmonton.ab.ca (accessed May 2001). 2. U.S. Environmental Protection Agency. In Biosolids Generation, Use, and Disposal in the United States, EPA530-R-99-009; September 1999. 3. Commission of the European Communities. Council directive on the protection of the environment, and in particular soil, when sewage sludge is used in agriculture (86=278=EEC). Official Journal of the European Communities 1986, No. 2, 181=6–12. 4. Maguire, R.O.; Sims, J.T.; Coale, F.J. Phosphorus solubility in biosolids-amended farm soils in the mid-atlantic region of the U.S. J. Environ. Qual. 2000, 29 (4), 1225–1233. 5. Sims, J.T.; Pierzynski, G.M. Assessing the impacts of agricultural, municipal, and industrial by-products on soil quality. In Land Application of Agricultural, Industrial, and Municipal By-Products; Power, J.F., Warren, A.D., Eds.; Soil Sci. Soc. Am.: Madison, WI, 2000; 237–261.

Fire and Soils P. K. Khanna Institute of Silviculture, University of Freiburg, Freiburg, Germany

R. J. Raison CSIRO Forestry and Forest Products, Kingston, Australian Capital Territory, Australia

INTRODUCTION

ATMOSPHERIC LOSSES OF ELEMENTS

Fires have played an integral role in the evolution and ecology of many ecosystems of the world. Natural and managed fires burn extensive areas annually,[1] and shape the landscape through effects on vegetation, soil, water, biodiversity, and socioeconomics. Fire has been actively managed for human benefit for millennia. In the tropics, slash-and-burn agriculture uses fire to release many of the nutrients accumulated in vegetation during a fallow period. Similarly, fire is widely used to remove forest and agricultural residues prior to establishment of a new crop and to enhance grazing. Fire is a dominant driver of the historical and current vegetation dynamics in forests, woodlands, shrublands, and grasslands in many parts of the world. Fire management practices (suppression, prescribed burning, etc.) remain controversial in many parts of the world because they can affect soil and a range of other factors with environmental values, including GHG emissions (CO2, CO, CH4, and NO2) and air quality.

During vegetation fires, atmospheric losses of elements occur in nonparticulate (e.g., C, N, S, and P) and particulate forms (Ca, Mg, K, and P), and postfire residues (partially burnt fuels, ash, and char components) are deposited on the soil. Depending upon the degree of combustion, a range of GHG gases are formed, the most important ones being CO2, CO, N2O, and CH4. Elements undergoing nonparticulate (gaseous) losses would undergo a long-range transport in the atmosphere (considered to be a complete loss to ecosystems), whereas particulates may be carried for only shorter distances. Raison et al.[4] used the conservation of Ca to distinguish between particulate and nonparticulate losses of different elements, and reported that a direct correlation between loss of N and loss of mass of plant material existed, providing the possibility of estimating accurately the loss of N from N content of the fuel and the amount of dry matter burnt. Up to 60% of P contained in the fuel may be lost, depending on factors such as the temperature, forms of P in the fuel, cation content of the ash, and the amount of ash transport. The nonparticulate transfer of P may range from 30% (low combustion) to 50% (high combustion). When combustion is relatively complete (gray ash is produced), nonparticulate losses of many elements may account for 60–80% of the total atmospheric transfer.[4]

FIRE CATEGORIES AND FIRE IMPACTS ON SOILS Fire can affect soil properties and processes both directly, as a result of combustion (via nutrient transfers to the atmosphere, ash, and char inputs, soil heating), and subsequently, as a result of a myriad of changes to ecosystem processes such as changed mineralization rates of soil organic matter and litter, erosion of ash and nutrient-rich surface soil, vegetation succession, and changes to N-fixing systems.[2] The scale of change ranges from minutes (soil heating) to many decades (vegetation succession), with the postfire impacts on soils often being the most dominant. Intensity and frequency of fires are major factors affecting the change in soils. The type and quantity of fuel consumed, and thus the duration of soil heating can be used to categorize fires in terms of impacts on soil and ecosystem processes (Table 1) 708 Copyright © 2006 by Taylor & Francis

EFFECTS OF SOIL HEATING Less than 10% of the heat produced during a vegetation fire is radiated downward, yet this heating is responsible for much of the direct changes in soil properties caused by forest fire.[5] The direct effects of heating on organic matter and soils range from mild sterilization and denaturing of protein at 50–60 C, to changes to clay minerals at 950 C.[3] Depending upon the degree of oxidation, carbonized products are produced, which range from mineral gray ash containing Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120025110 Copyright # 2006 by Taylor & Francis. All rights reserved.

Fire and Soils

709

Table 1 Fire categories as defined by fuel type, fuel quantity, and the duration of soil heating

Fire category

Fuel types

Fuel consumption (t/ha)

Maximum temperature ( C) at 2-cm depth

Time (hr) for which soil is heated above 80% of the maximum temperature

Frequency of occurrence (yr)

Relative impact on soils

Windrows

Mostly trunks and slash

100–400þ

500

>24

30–100þ

High

Forest regeneration=slash

Tree crowns, some logs, shrubs, and litter

50–300þ

200

0.5–2

30–100þ

High

Forest wildfire

Litter, shrubs, and crowns

20–60

100

0.1–1

20–100þ

Medium–high

Shrubland

Litter and crowns

15–30

80

0.1–0.5

20þ

Medium–low

Grassland, crop residue

Litter, crop residues

0.5–5.0

60

0.1–0.3

1þ

Low

Values are a guide, with considerable variation observed. (From Ref.[3].)

very little residual C to black residue with a high amount of C and charred substances. Moderate heating can render soil organic matter more prone to microbial respiration under postfire conditions.[3] Soil heating may be as important as the addition of ash in slash-and-burn systems, and can increase mineral N and P fractions.[6]

EFFECTS OF ASH ADDITIONS After fire, postfire residues deposit highly aromatic (char) carbon. About 1–5% of burnt fuel is converted to char, and this is considered to be relatively inert and likely to be long-lived in soils[7] and sediments. Elements, except C and N, are enriched when ash is formed, and the level of their enrichment depends on the degree of combustion and initial fuel characteristics. In comparison with unburnt eucalypt fuels, concentrations of Ca, Mg, and P increased by 10- to 50-fold, 10- to 35-fold, and 10-fold, respectively,[8] with high values in gray ash. As ash is prone to be transported by wind and water, any loss of ash can cause major losses of elements from ecosystems. Ash is highly alkaline and will increase the pH of the soil and its capacity to buffer protons. The salt content of ash may initially inhibit seedling growth, but plants growing around an ashbed benefit from the nutrients made available in the short- and long term, causing the so-called ‘‘ashbed effect.’’ Ash can also stimulate the soil biological activity[9] and interact positively with the effects of soil heating.[10]


CHANGES IN SOIL PHYSICAL PROCESSES Fire affects water penetration and erosion susceptibility, by creating hydrophobicity[11] in the surface soil. Erosion has major effects on soil fertility. Blackening of the soil surface and removal of vegetation and litter can increase soil temperature. Where fire reduces the vegetation cover, frost damage may increase, plant water uptake is reduced, thereby increasing the soil wetness, and, sometimes, rates of mineralization of soil organic matter and leaching of nutrients.

CHANGES IN SOIL CHEMICAL PROPERTIES AND PROCESSES Large quantities of cations (Ca, Mg, K, and NH4) and anions (Cl and SO4) and soluble silica are mobilized in surface soils by vegetation fire, especially under ashbeds[12] where Ca remained the main cation in the solution phase of surface soil during a three-year study period. Increase in exchange capacity (because of the pH change in acid soils) and the exchangeable base cations occurs after fire and may remain so for many years after intense fire. Nitrification rates and hence the potential for leaching of nitrate and cations may be increased after intense fire.[2] A small fraction (about 10%) of total P in ash may be soluble in water.[9] Protons are needed to mobilize P deposited in ash and, therefore, mixing of ash with acid soil is required before P can be used by postburn vegetation. The P-sorption capacity of soils usually increases after intense forest fire.

710

CHANGES IN BIOLOGICAL PROPERTIES OF SOILS The effects of fire on biological processes are complex and can be very site dependent. Fire often enhances the decomposition of organic matter and the mineralization of organically bound elements. This can be because of the changed microbial populations, altered substrate quality (additional soluble C), and more favorable environmental factors (temperature, moisture, and pH). For example, in laboratory experiments where ash was added to different soils, enhanced respiration rates, especially in soils with high organic matter content, were observed during 6 weeks of incubation[9] until the easily mineralizable soil C was respired.[13] Nitrogen mineralization rates and the amount of nitrate produced usually increase following a fire, and in some cases may even lead to denitrification losses.[9] A change in microbial population (autotrophic nitrifiers replacing heterotrophic nitrifiers) after fire has been proposed by Bauhus, Khanna, and Raison.[13] Study of denitrification rates after fire deserves greater research attention. MANAGEMENT OF SOILS FOLLOWING FIRE One of the immediate concerns following vegetation fires is to minimize soil erosion and to replenish nutrients that are lost by atmospheric transfer or erosion. Losses of N can be large (several hundred kg=ha) in forests that are subjected to wildfires or slash burns. In natural systems, given sufficient time, N-fixing processes can maintain N balance. Replenishment of lost P and cations may take decades or centuries, which may affect plant productivity if such fires occur frequently. In managed systems, fertilizer inputs may be needed to maintain productivity. REFERENCES 1. UNEP. Wild Fires, a Double Impact on the Planet, 2005. http:==www.grid.unep.ch=product=publication= earlywarning.php.


Fire and Soils

2. Raison, R.J.; O’Connell, A.M.; Khanna, P.K.; Keith, H. Effects of repeated fires on nitrogen and phosphorus budgets and cycling processes in forest ecosystems. In Fire in Mediterranean Ecosystems, Report No. 5, Ecosystem Research Series; Trabaud, L., Prodon, R., Eds.; Commission of the European Communities: Brussels, 1993; 347–363. 3. Walker, J.; Raison, R.J.; Khanna, P.K. Fire. In Australian Soils—The Human Impact; Russell, J.S., Isbell, R.F., Eds.; Queensland University Press: Brisbane, 1986; 186–216. 4. Raison, R.J.; Khanna, P.K.; Woods, P.V. Mechanisms of element transfer to the atmosphere during vegetation fires. Can. J. For. Res. 1985, 15, 132–140. 5. Neary, D.G.; Klopatek, C.C.; DeBano, L.F.; Folliott, P.F. Fire effects on belowground sustainability: a review and synthesis. For. Ecol. Manage. 1999, 122, 51–71. 6. Giardina, C.P.; Sanford, R.L.; Dockersmith, I.C. Changes in soil phosphorus and nitrogen during slashand-burn clearing of a dry tropical forests. Soil Sci. Soc. Am. J. 2000, 64, 399–405. 7. Skjemstad, J.O.; Taylor, J.A.; Smernik, R.J. Estimation of charcoal (char) in soil. Commun. Soil Sci. Plant Annal. 1999, 30, 2283–2298. 8. Raison, R.J.; Khanna, P.K.; Woods, P.V. Transfer of elements to the atmosphere during low-intensity prescribed fires in three Australian subalpine eucalypt forests. Can. J. For. Res. 1985, 15, 657–664. 9. Khanna, P.K.; Raison, R.J.; Falkiner, R.A. Chemical properties of ash components derived from Eucalyptus litter and its effects on forest soils. For. Ecol. Manage. 1994, 66, 107–125. 10. Raison, R.J. Modification of the soil environment by vegetation fires, with particular reference to nitrogen transformations: a review. Plant Soil 1979, 51, 73–108. 11. DeBano, L.F.; Neary, D.G.; Folliott, P.F. Fire’s Effects on Ecosystems; John Wiley & Sons: New York, U.S.A., 1998. 12. Khanna, P.K.; Raison, R.J. Effects of fire intensity on solution chemistry of surface soil under a Eucalyptus pauciflora forest. Aust. J. Soil Res. 1986, 24, 423–434. 13. Bauhus, J.; Khanna, P.K.; Raison, R.J. The effect of fire on carbon and nitrogen mineralization and nitrification in an Australian forest soil. Aust. J. Soil Res. 1993, 31, 621–639.

Flooding Tolerance of Crops Tara T. VanToai Getachew Boru Jianhuan Zhang United States Department of Agriculture (USDA), The Ohio State University, Columbus, Ohio, U.S.A.

INTRODUCTION Periodic flooding during the growing season adversely affects crop growth and productivity, with the exception of flooded rice, in many areas of the U.S. and the rest of the world. Soil can become flooded when it is poorly drained or when rainfall or irrigation is excessive. Other terms, such as soil saturation, waterlogging, anoxia, and hypoxia, are also commonly used to describe flooding conditions. Flooding causes premature senescence, which results in leaf chlorosis, necrosis, defoliation, reduced nitrogen fixation, cessation of growth, and reduced yield. The severity of the flooding stress is affected by many factors, including flooding duration, crop variety, growth stage, soil type, fertility levels, pathogens, and flooding conditions.[1] In general, stream flooding, characterized by the overflow of rivers or creeks into a flood plain, is more damaging than lowland flooding, characterized by inadequate surface drainage and slow soil permeability of depressional areas. Sediments carried by stream flooding, when deposited on the leaves of flooded plants, can cause severe wilting and plant death within 24 hr of the stress. Flooding can be further divided into either waterlogging, where only the roots are flooded, or complete submergence, where the entire plants are under water. While plants develop adaptive mechanisms to allow them to survive long-term waterlogging, most plants die within one or two days of submergence.[1] The lack of oxygen has been proposed as the main problem associated with flooding.[2] Indeed, tolerance of anoxia and hypoxia has been used synonymously with tolerance of flooding stress. During the last two decades, a great deal more information has accumulated from research on the molecular, biochemical, and physiological responses of plants to the lack of oxygen rather than to flooding per se.[3–5] However, tolerance of field flooding appears to be much more complex than tolerance of artificially induced hypoxia and anoxia. Contrary to the injury seen in flooded fields, soybeans can thrive in stagnant water in the greenhouse and soybeans grown in hydroponic medium continuously bubbled with nitrogen gas, where the dissolved oxygen level was not detectable, showed no symptoms of stress.[6] Soybean, therefore, is much Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001680 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

more tolerant to excessive water and a lack of oxygen than previously expected. The reasons underlying the dramatic differences between responses to flooding in the greenhouse and flooding in the field are not known. However, growth reduction and yield loss in flooded fields could arise from root rot diseases, nitrogen deficiency, nutrient imbalance, and=or the accumulation of toxic levels of CO2 in the root zone. Indeed, the levels of CO2 commonly found in flooded soil (30%) severely caused leaf chlorosis and reduced plant biomass of soybean, a flood-susceptible crop, but not of rice, a flooding-tolerant crop (VanToai, unreported data).

MORPHOLOGICAL AND ANATOMICAL ADAPTATION TO FLOODING STRESS One important morphological change associated with flooded roots is the formation of aerenchyma tissue, which contains continuous gas-filled channels connecting the root with the shoot. Other morphological changes, including hypertrophy and the formation of lenticels, adventitious roots, and pneumatophores, have also been observed in many plant species.[7] Flooding can also change the direction of root growth. Roots of tomato and sunflower plants become disgeotropic or negatively geotropic under flooding conditions instead of positively geotropic.[7] The changes in orientation of roots under flooding conditions enables them to escape stress from the reduced oxygen availability by growing closer to the better aerated soil surface.

PHYSIOLOGICAL, BIOCHEMICAL, AND MOLECULAR ADAPTATION TO FLOODING STRESS Rice cultivars that showed rapid leaf and sheath growth during submergence did not survive as well as cultivars that did not elongate. Tolerant cultivars appeared to conserve carbohydrates in the shoots and roots during submergence.[8] Upon removal of the stress, tolerant rice cultivars are able to recover more rapidly and suffer less plant mortality.[9] An adequate 711

712

supply of sugar is needed for corn root tips to survive anoxic and hypoxic stresses.[10] The lack of oxygen induces a set of anaerobic proteins in roots to allow plants to cope with the stress.[11] These stress proteins are enzymes of either glycolysis, glucose metabolism, or fermentation.[10,12]

GENETIC VARIABILITY IN FLOODING TOLERANCE Flooding tolerance is usually defined as minimal or no yield loss. According to VanToai et al.,[13] waterlogging for four weeks during the early flowering stage reduced the average grain yield of 84 U.S. soybean cultivars by 25%. Yield reduction, however, varied from 9% in the most flooding tolerant cultivar to 75% in the most flooding susceptible cultivar. Flooding tolerance can also be defined as high yield under flooding stress. According to this definition, the most flooding tolerant variety in this study produced 3.7 Mg ha1, while the least produced 1.27 Mg ha1. When the cultivars were ranked for flooding tolerance based on both definitions, seven of the ten most flooding tolerant cultivars were the same; and seven of the ten least flooding tolerant cultivars were also the same. Thus, the two definitions of flooding tolerance, either high yield under flooding or minimal yield difference between nonflooded and flooded conditions, appear to be compatible. Flooding tolerance is independent from nonflooded yield indicating that genetic variability for flooding tolerance exists and could be improved through plant breeding and selection. Studies of submergence tolerance of rice showed that the unimproved land races FR13A, Janki, and FR43B had survival values ranging from 41 to 51% after 10 days of submergence, while only 2–4% of elite cultivars (IR74, IR48 and IR68) survived.[9]

IMPROVING FLOODING TOLERANCE BY TRADITIONAL PLANT BREEDING Tolerance of flooding in wheat (Triticum aestivum L.) and rice (Oryza sativa L.) is a quantitative trait controlled by a small number of genes.[9,14] Using the submergence tolerant land races FR13A, Janki, and FR43B as donor parents, rice breeders at the International Rice Research Institute (IRRI) at Los Banos in the Philippines have developed an experimental rice line (IR49830-7-1-2-2) from crosses with the short stature, high-yield IR lines, which produced as much as 4880 kg ha1.[9] The result showed that submergence tolerance can be incorporated into improved, highyielding cultivars to raise the productivity in submergence-prone areas.


Flooding Tolerance of Crops

IDENTIFYING FLOODING TOLERANCE LOCI AND IMPROVING FLOODING TOLERANCE BY MOLECULAR PLANT BREEDING Xu and Mackill[15] identified a single submergence tolerance locus, Sub1, that controls about 50% of the variation in submergence tolerance of rice. During the last few years, molecular marker aided selection has been used successfully for the breeding of crops with improved quantitative traits. VanToai et al.[16] identified a single DNA marker that was associated with improved plant growth (from 11 to 18%) and grain yields (from 47 to 180%) of soybean in waterlogged environments. The identified marker was uniquely associated with waterlogging tolerance and wasnot associated with maturity, normal plant height or grain yield. Near isogenic lines with and without the flooding tolerant marker have been developed and are being field tested under waterlogging conditions to confirm the association of the marker with the tolerance of soybean to waterlogging stress.

IMPROVING FLOODING TOLERANCE BY GENETIC TRANSFORMATION Flooding induces or accelerates plant senescence in tobacco, tomato, sunflower, carrot, barley, peas, wheat, maize, and soybean. The most obvious visual symptom of flooded plants under stress is the yellowing of leaves followed by necrosis due to premature senescence. Within one day of flooding, the concentration of the antiaging hormone, cytokinin, in sunflower xylem sap declined sharply to a very low level.[17] In order to test if enhanced endogenous cytokinin production could improve flooding tolerance, Zhang et al.[18] generated transgenic plants containing a gene coding for cytokinin biosynthesis. Four transgenic Arabidopsis lines were chosen for cytokinin and flooding tolerance determinations. The levels of cytokinin were similar between wild-type and transgenic plants in the unflooded treatment. After 5 days of waterlogging, the cytokinin increased 3–10 times in transgenic plants as compared to wild-type plants. In three independent experiments, all four transgenic lines were consistently more tolerant to soil waterlogging and complete submergence than wild-type plants. The results indicated that endogenously produced cytokinin can regulate senescence caused by flooding stress, thereby increasing plant tolerance of flooding. This study provides a novel mechanism to improve flooding tolerance in plants.[18] In summary, while the lack of oxygen has been used interchangeably with flooding stress, tolerance of field flooding is more complex than tolerance of anoxia and hypoxia. The use of molecular plant breeding

Flooding Tolerance of Crops

and genomic transformation to improve flooding tolerance in crops is promising to be successful.

REFERENCES 1. Sullivan, M.; VanToai, T.; Fausey, N.; Beuerlein, J.; Parkinson, R.; Soboyejo, A. Evaluating on-farm flooding impacts on soybean. Crop Science 2001, 41, 1–8. 2. Kozlowski, T.T. Extent, causes, and impacts of flooding. In Flooding and Plant Growth; Kozlowski, T.T., Ed.; Academic Press Inc.: Orlando, FL, 1984; 1–7. 3. Kennedy, R.A.; Rumpho, M.E.; Fox, T.C. Anaerobic metabolism in plants. Plant Physiology 1992, 84, 1204–1209. 4. Perata, V.M.; Alpi, A. Plant responses to anaerobiosis. Plant Science 1993, 93, 1–17. 5. Ricard, B.; Couee, I.; Raymond, P.; Saglio, P.H.; Saint-Ges, V.; Pradet, A. Plant metabolism under hypoxia and anoxia. Plant Biochemistry 1994, 32, 1–10. 6. Boru, G.; VanToai, T.; Alves, J.D. Flooding injuries in soybean are caused by elevated carbon dioxide levels in the root zone. Fifth National Symposium on Stand Establishment 1997, 205–209. 7. Hook, D.D. Adaptations to flooding with fresh water. In Flooding and Plant Growth; Kolowski, T.T., Ed.; Academic Press: Orlando, FL, 1984; 265–294. 8. Jackson, M.B.; Waters, I.; Setter, T.; Greenway, H. Injury to rice plants caused by complete submergence: a contribution by ethylene (ethene). Journal of Experimental Botany 1987, 38, 1826–1838.


713

9. Mackill, D.J.; Amante, M.M.; Vergara, B.S.; Sarkarung, S. Improved semidwarf rice lines with tolerance to submergence of seedlings. Crop Science 1993, 33, 749–753. 10. Ricard, B.; Saglio, P.; VanToai, T.; Chourey, P. Evidence for the critical role of sucrose synthase for anoxic tolerance of maize roots using a double mutant. Plant Physiology 1998, 116, 1323–1331. 11. Sachs, M.M.; Freeling, M.; Okimoto, R. The anaerobic proteins of maize. Cell 1980, 20, 761–767. 12. Sachs, M.M.; Subbaiah, C.C.; Sabb, I.N. Anaerobic gene expression and flooding tolerance in maize. Journal of Experimental Botany 1996, 47, 1–15. 13. VanToai, T.; Beuerlein, J.E.; Schmitthenner, A.F.; St. Martin, S.K. Genetic variability for flooding tolerance in soybeans. Crop Science 1994, 34, 1112–1115. 14. Boru, G.; Ginkel, V.M.; Kronstad, W.E.; Boersma, L. Expression and inheritance of tolerance to waterlogging stress in wheat. Euphytica 2001, 117, 91–98. 15. Xu, K.; Mackill, D.J. A major locus for submergence tolerance mapped on rice chromosome 9. Molecular Breeding 1996, 2, 219–224. 16. VanToai, T.T.; St. Martin, S.K.; Chase, K.; Boru, G.; Schnipke, V.; Schmitthenner, A.F.; Lark, K.G. Identification of a QTL associated with tolerance of soybean to soil waterlogging. Crop Science 2001, 41, 1247–1252. 17. Burrows, W.J.; Carr, D.J. Effects of flooding the root system of sunflower plants on the cytokinin content of the xylem sap. Physiol Plant. 1969, 22, 1105–1112. 18. Zhang, J.; VanToai, T.; Huynh, L.; Preiszner, J. Development of flooding-tolerant Arabidopsis Thaliana by autoregulated cytokinin production. Molecular Breeding 2000, 6, 135–144.

Fluid Flow: Challenges Modeling John L. Nieber University of Minnesota, St. Paul, Minnesota, U.S.A.

INTRODUCTION The unsaturated zone (vadose zone) of the Earth’s crust is an important interface for both the underlying groundwater and the overlying surface water resources and atmosphere. Quantifying fluid flow, mass transport, and energy transport processes in the unsaturated zone have become a focus for researchers, government agencies, and consultants during the past three decades because it is found that the outcome of these processes have an impact on the sustainability of modern social structures. While sophisticated sensors and instrumentation seems to have been developed to provide data from the field on a real-time basis, these cannot be used to directly predict the possible outcomes. Instead, this prediction requires the use of mathematical models representing the flow and transport processes. The following describes some of the applications of models for flow in the unsaturated zone, past challenges and achievements in improving modeling methods, and future challenges to modeling flow.

APPLICATION OF MODELS Models for simulating flow processes in soil and groundwater find applications in many environmentally oriented disciplines including geography, soil science, agricultural engineering, civil engineering, geoengineering, hydrogeology, and meteorology. These models are generally based on numerical solutions of governing equations and require the use of digital computers to complete the calculation task. There are many current and potential applications for such models, but a short list of the common applications includes: 1) estimation of groundwater recharge volume and contaminant loading; 2) design of measures to remediate contaminated soils and groundwater; 3) design of efficient drainage and irrigation systems for efficient crop production; 4) estimation of runoff production from land in response to rainfall and/or snowmelt; and 5) assessment of the impact of global climate change on surface and subsurface water resources. 714 Copyright © 2006 by Taylor & Francis

Models developed for these applications need to satisfy two criteria to be put to use by a practitioner. They have to be fairly easy to use and dependable. The first criterion, ease-of-use, only requires a good team of programmers, and does not pose a challenge for flow modeling. The second criterion stipulates that the model provides accurate results in a timely manner. This is a direct challenge to flow modeling. The following sections present information on the past and future challenges associated with the development of dependable models.

PAST CHALLENGES AND ACHIEVEMENTS During the past three decades there has been substantial progress in the development of numerical models for simulating relevant unsaturated zone processes. During this time predominant attention in modeling flow processes was given to the Richards equation.[1] In the past, many of the difficulties in model development involved the determination of the best ways to solve this equation. Due to the highly nonlinear character of the Richards equation, analytical solutions [2] to the equation were possible only for simplified conditions, and therefore numerical solution methods were necessary to treat realistic field conditions. Numerical methods such as the finite difference method and the finite element method were adopted from other engineering and science applications and applied to discretize the Richards equation into systems of nonlinear algebraic equations. Two major numerical solution problems were faced by researchers in solving this equation. One was the highly nonlinear nature of the equation and associated boundary conditions, leading to problems of slow convergence or even nonconvergence of the solution methods. This problem was handled by applying nonlinear equation solvers classified within the broad class of Newton methods and Picard methods.[3] A second problem was the need to be able to solve large systems of algebraic equations, which in earlier years involved hundreds or perhaps thousands of equations. Due to the relatively small memory Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001757 Copyright # 2006 by Taylor & Francis. All rights reserved.

Fluid Flow: Challenges Modeling

available on early computers it was necessary to use iterative methods to solve even moderate-sized problems. Conventional iterative methods like Jacobi, Gauss–Seidel, and successive over-relaxation methods[4] were used to solve these problems. While these methods were satisfactory for relatively homogeneous systems, for strongly heterogeneous systems it was found that these methods led to poor convergence or even nonconvergence. Computer memory did increase substantially during the 1980s and 1990s, and this allowed the use of direct equation solvers for the types of problems solved in the earlier years. However, with the increase of computer memory storage, the problems tackled have also increased memory requirements (tens of thousands to millions of equations need to be solved) and iterative methods are once again back in vogue. Fortunately, the efficiency of iterative methods has also increased substantially with the development of conjugate gradient methods[5] and multigrid methods.[6] During the last 30 years there has been an ongoing effort to find the recipe for a numerical method or set of methods that will be robust for solving the Richards equation.[3,4,7–12] The desire has been to develop computationally efficient methods applicable to a broad range of practical problems, especially for large-scale, three-dimensional, heterogeneous flow systems. Associated problems involved assuring that the solutions were mass conservative and that the iterative methods used to solve the nonlinear algebraic equations would converge even for conditions where the soil is very dry and highly heterogeneous. Detailed analysis of how to assure a mass conservative solution has been given in Ref.[8]. Assurance of nonlinear iteration convergence has been found to be more problematic and recent improvements have been made using techniques involving primary variable switching,[9,10] higher order time integration,[11] and variable transformation.[12] Aside from the obvious problems associated with solving the Richards equation there has been the need to assign (spatially) the equation parameters for field scale applications of the numerical solutions. For instance, the catchment scale model of Ref.[13], like the model of Ref.[4], was developed to facilitate the simulation of three-dimensional variably saturated flow over an entire catchment. While the solution of the large system of algebraic equations for such a problem offers a significant challenge to modeling, an equally if not larger challenge is the problem of assigning parameters to the cells in the numerical grid. The problem of parameter assignment is two-fold. For one, there is the problem of determining the spatial distribution of the parameters at the scale of the numerical grid cell. Involved with this problem is the uncertainty in predicting these values, given limited field data. Geostatistical methods[14] were developed


715

to facilitate such prediction and to estimate limits of uncertainty. A second problem arises because small scale features, and even small scale processes, occurring below the sampling scale of the field scale grid cell can significantly influence the actual outcome at the field scale, but are not directly predictable by the governing equation(s) discretized at the field scale grid. There has been some success[15] in developing techniques that provide parameters that account for these subgrid scale features and processes. These parameters, called effective parameters, are those that yield effectively the same outcome as would occur if the more detailed parameter distribution were used.

FUTURE CHALLENGES TO MODELING Two of the greatest future challenges to modeling include the need to more completely describe the flow and transport processes, and the need to incorporate multiscale phenomena into the modeling analysis. The first challenge involves the expansion of the governing equations to include coupled physical and chemical processes. The second challenge involves the assignment of equation parameters and the incorporation of subgrid features and processes. To assure the success of future modeling, this second challenge is the most critical to address.

Governing Equations As awareness of environmental problems has increased, and environmental regulations have become more stringent, the scope for modeling has expanded from modeling the flow of water alone to modeling coupled multiphase fluid flows, mass (solute) transport and energy (thermal) transport. The coupled equations for isothermal multiphase fluid flow have been reviewed in detail.[16] Numerical solutions of coupled two-phase flow equations for applications to environmental problems have advanced considerably in the last 20 years. One of the earlier multiphase flow solutions is given in Ref.[17], while methods representing the latest advances in computational efficiency are given in Ref.[18]. Equations for nonisothermal conditions, which until recently have received much less attention, have been presented in Ref.[19]. Solutions of these equations offer new challenges to those developing numerical solutions.[20,21] Recent techniques such as those mentioned before for the solution of the Richards equation and the coupled multiphase flow equations

716

should be tested to improve the efficiency of such solutions.

Process Scale Numerical solutions of flow and transport processes for field-scale applications are generally performed on relatively large numerical grids. The reason for this is two-fold. First, the computational effort to simulate field-scale problems can easily exceed the capabilities of today’s computers (using current numerical methods) if too fine a grid is used. Second, one generally does not know what values of parameters to assign to very fine grids due to inadequate spatial resolution in field data. As a result, it is necessary to assign effective parameters to the field-scale grid cells. One approach is to use stochastic methods[15] to derive effective parameters. Renormalization methods,[22] which rely on numerical methods, are another means to derive effective parameters. In some instances the governing equations may behave differently at the small scale than at the large scale. For instance, in the case of finger flow[23,24] or funnel flow,[25] the flow occurs on a scale of about 10 cm. To simulate these flow features directly it is necessary to use relatively small grid cells. Using large grid cells without considering these small-scale processes leads to a diffuse solution and the small-scale features are missed. Effectively capturing the dynamics of these small-scale processes into a field-scale grid requires an appropriate procedure for upscaling of the governing equations from the small-scale to the field-grid scale. Such a procedure has been demonstrated[26] for viscous fingering in multiphase flow and is an area of active research. The discussion about process scale poses the question about whether the governing equations maintain a constant form in the progression from the small scale to the large scale. Addressing this question involves the principles currently developed in the field of multiscale science.[27] While these developments originated in the fields of solid mechanics and fluid mechanics, they are currently receiving much attention in hydrology, soil physics, and hydrogeology.

REFERENCES 1. Richards, L.A. Capillary conduction of liquids through porous mediums. Physics 1931, 1, 318–333. 2. Philip, J.R. Steady infiltration from spheroidal cavities in isotropic and anisotropic soils. Water Resour. Res. 1986, 22, 1874–1880.



3. Paniconi, C.; Putti, M.A. A comparison of picard and newton iteration in the numerical solution of multidimensional variably saturated flow problems. Water Resour. Res. 1994, 30, 3357–3374. 4. Freeze, R.A. Three-dimensional, transient, saturatedunsaturated flow in a groundwater basin. Water Resour. Res. 1971, 7, 347–366. 5. Saad, Y. Iterative Methods for Sparse Linear Systems; PWS Publ. Co.: Boston, MA, 1996. 6. Brandt, A. Multi-level adaptive solutions to boundary value problems. Math. Comput. 1977, 31, 333–390. 7. Huyakorn, P.S.; Thomas, S.D.; Thompson, B.M. Techniques for making finite elements competitive in modeling flow in variably saturated media. Water Resour. Res. 1986, 20, 1099–1115. 8. Rathfelder, K.; Abriola, L.M. Mass conservative numerical solutions of the head-based richards equation. Water Resour. Res. 1994, 30, 2579–2586. 9. Forsyth, P.A.; Wu, Y.S.; Pruess, K. Robust numerical methods for saturated-unsaturated flow with dry initial conditions in heterogeneous media. Adv. Water Resour. 1995, 18, 844–856. 10. Diersch, H.-J.G.; Perrochet, P. On the primary variable switching technique for simulating unsaturatedsaturated flows. Adv. Water Resour. 1999, 23, 271–301. 11. Tocci, M.D.; Kelley, C.T.; Miller, C.T. Accurate and economical solution of the pressure-head form of richards’ equation by the method of lines. Adv. Water Resour. 1997, 20, 1–14. 12. Williams, G.A.; Miller, C.T. An evaluation of temporally adaptive transformation approaches for solving richards’ equation. Adv. Water Resour. 1999, 22, 831–840. 13. Paniconi, C.; Wood, E.F. A detailed model for simulation of catchment scale subsurface hydrologic processes. Water Resour. Res. 1993, 29, 1601–1620. 14. Webster, R. Quantitative spatial analysis of soil in the field. In Advances in Soil Science; Stewart, B.A., Ed.; Springer Verlag: New York, 1985; Vol. 3, 1–70. 15. Mantoglou, A.; Gelhar, L.W. Stochastic modeling of large-scale transient unsaturated flow systems. Water Resour. Res. 1987, 23, 37–46. 16. Miller, C.T.; Christakos, G.; Imhoff, P.T.; McBride, J.F.; Pedit, J.; Trangenstein, J.A. Multiphase flow and transport modeling in heterogeneous porous media: challenges and approaches. Adv. Water Resour. 1999, 21, 77–120. 17. Kaluarachchi, J.J.; Parker, J.C. An efficient finite element method for modeling multiphase flow. Water Resour. Res. 1989, 25, 43–54. 18. Bastian, P.; Helmig, R. Efficient fully-coupled solution techniques for two-phase flow in porous media; parallel multigrid solution and large scale computations. Adv. Water Resour. 1999, 23, 199–216. 19. Nassar, I.N.; Horton, R. Heat, water and solute transfer in unsaturated porous media: I. Theory development and transport coefficient evaluation. Trans. Por. Med. 1997, 27, 17–39. 20. Thomas, H.R.; Missoum, H. Three-dimensional coupled heat, moisture, and air transfer in a deformable


unsaturated soil. Int. J. Num. Meth. Engng. 1999, 44, 919–943. 21. Chounet, L.M.; Hilhorst, D.; Jouron, C.; Kelanemer, Y.; Nicolas, P. Simulation of water flow and heat transfer in soils by means of a mixed finite element method. Adv. Water Resour. 1999, 22, 445–460. 22. King, P.R. The use of renormalization for calculating effective permeability. Transp. Por. Med. 1989, 4, 37–58. 23. Glass, R.J.; Steenhuis, T.S.; Parlange, J.-Y. Wetting front instability as a rapid and far-reaching hydrologic process in the vadose zone. J. Contam. Hydrol. 1988, 3, 207–226.


717

24. Nieber, J.L. Modeling finger development and persistence in initially dry porous media. Geoderma 1996, 70, 209–229. 25. Ju, S.-H.; Kung, K.-J.S. Steady-state funnel flow: its characteristics and impact on modeling. Soil Sci. Soc. Am. J. 1997, 61, 416–427. 26. Blunt, M.; Christie, M. How to predict viscous fingering in three-component flow. Transp. Por. Med. 1993, 12, 207–236. 27. Glimm, J.; Sharp, D.H. Multiscale science: a challenge for the twenty-first century. SIAM News 1997, 304, 17, 19.

Forest Ecosystems: Nutrient Cycling Neil W. Foster Canadian Forest Service, Ontario, Sault Ste. Marie, Ontario, Canada

Jagtar S. Bhatti Canadian Forest Service, Northern Forestry Centre, Edmonton, Alberta, Canada

INTRODUCTION

INFLUENCE OF CLIMATE

Nutrients are elements or compounds that are essential for the growth and survival of plants. Plants require large amounts of nutrients such as nitrogen (N), phosphorus (P), carbon (C), hydrogen (H), oxygen (O), potassium (K), calcium (Ca), and magnesium (Mg), but only small amounts of others such as boron (B), manganese (Mn), iron (Fe), copper (Cu), zinc (Zn) and chlorine (Cl) (micronutrients). Forest nutrient cycling is defined as the exchange of elements between the living and nonliving components of an ecosystem.[1] The processes of the forest nutrient cycle include: nutrient uptake and storage in vegetation perennial tissues, litter production, litter decomposition, nutrient transformations by soil fauna and flora, nutrient inputs from the atmosphere and the weathering of primary minerals, and nutrient export from the soil by leaching and gaseous transfers. Each nutrient element is characterized by a unique biogeochemical cycle. Some of the key features of the major nutrients are shown in Table 1. Forest trees make less demand on the soil for nutrients than annual crops because a large proportion of absorbed nutrients are returned annually to the soil in leaf and fine root litter and are reabsorbed after biological breakdown of litter materials. Also, a large portion of nutrient requirement of trees are met through internal cycling as compared with agricultural crops. Nutrient cycling in forest ecosystems is controlled primarily by three key factors: climate, site, abiotic properties (topography, parent material), and biotic communities. The role of each factor in ecosystem nutrient dynamics is discussed and illustrated with selected examples from boreal, temperate, and tropical zones. The importance of ecosystem disturbance to nutrient cycling is examined briefly, since some nutrients are added or lost from forest ecosystems through natural (e.g., fire, erosion, leaching) or human activity (harvesting, fertilization).

Large-scale patterns in terrestrial primary productivity have been explained by climatic variables. In aboveground vegetation, nutrient storage generally increases in the order: boreal < temperate < tropical forests (Table 2). In contrast, forest floor nutrient content and residence time increases from tropical to boreal forests, as a result of slower decomposition in the cold conditions of higher latitudes. In subarctic woodland soils and Alaskan taiga forests, nutrient cycling rates are low because of extreme environmental conditions.[2] Arctic and subarctic forest ecosystems have lower rates of nutrient turnover and primary production because of low soil temperature, a short growing season, low net AET and the occurrence of permafrost. Low temperature reduces microbial activity, litter decomposition rates, and nutrient availability and increases C accumulation in soil. In contrast with high latitudes, conditions in a tropical forest favor microbial activity throughout the year, which generally results in faster decomposition except in situations with periodic flooding, soil dessication, and low litter quality.[3] Rates of plant material decay are an order of magnitude higher in tropical soils than in subarctic woodland soils. The low storage of C and high amount of litter production in highly productive tropical forests contrasts with the high C storage and low litter production in boreal forests (Table 2).

718 Copyright © 2006 by Taylor & Francis

INFLUENCE OF BIOTIC FACTORS Nutrient cycles are modified substantially by tree species-specific controls over resource use efficiency (nutrient use per unit net primary production). Species vary widely in their inherent nutrient requirements and use.[4] These effects can be split into two categories: accumulation into living phytomass and production of various types of nutrient-containing dead Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001709 Copyright # 2006 by Taylor & Francis. All rights reserved.

Forest Ecosystems: Nutrient Cycling

719

Table 1 Features of the major nutrient cycles Element

Major sources for tree uptake

Uptake by the trees

Limiting situations

Carbon

Atmosphere

Atmosphere

Atmospheric concentration may limit growth

Oxygen

Atmosphere

Atmosphere

Waterlogged soils

Hydrogen

Atmosphere

Atmosphere

Extremely acidic and alkaline conditions

Nitrogen

Soluble NO3 and NH4; N2 for nitrogen fixing species

Soil organic matter; atmospheric N2 for nitrogen fixing species

Most temperate forests, many boreal forests and some tropical forests

Phosphorus

Soluble phosphorus

Soil organic matter; adsorbed phosphate and mineral phosphorus

Old soils high in iron and aluminum, common in subtropical and tropical environment

Potassium calcium magnesium

Soluble Kþ=soluble Ca2þ= soluble Mg2þ

Soil organic matter; exchange complex and minerals

Miscellaneous situations and some old soils

phytomass. Rapid accumulation of phytomass is associated with a net movement of nutrients from soil into vegetation. More than half of the annual nutrient uptake by a forest is typically returned to forest floor (litterfall) and soil (fine-root turnover). The subsequent recycling of these nutrients is a major source of available nutrients for forest vegetation. The mean annual litterfall from above-ground vegetation increases from boreal regions to the tropics following the gradient of productivity (Table 2). Nutrient availability is strongly influenced by the quantity and quality of litter produced in a forest. A high proportion of the variation in foliar N concentrations at the continental scale has been

explained by differences between forest types, which in turn has large impact on litter quality and the nutrient content of forest floors. In many temperate and boreal forest ecosystems, microbial requirement for N increases or decreases with labile supplies of soil C. Increased microbial demand for N may temporarily decrease the N availability to trees during the initial decomposition of forest residues with a wide C/N ratio. Microbes immobilize N from the surrounding soil, relatively rapid for readily decomposable organic matter (needle litter), and more slowly for recalcitrant material (branches, boles). Rates of net N mineralization are higher and retention of foliar N is lower in temperate and tropical than

Table 2 Nutrient distribution in different forest ecosystems Vegetation (Mg ha1)

Forest floor (Mg ha1)

Soil (Mg ha1)

Residence time (year)

Carbon Boreal coniferous Temperate deciduous Tropical rain forest

78–93 103–367 332–359

37–113 42–105 7–72

41–207 185–223 2–188

800 200 120

Nitrogen Boreal coniferous Temperate deciduous Tropical rain forest

0.3–0.5 0.1–1.2 1.0–4.0

0.6–1.1 0.2–1.0 0.03–0.05

0.7–2.87 2.0–9.45 5.0–19.2

200 6 0.6

Phosphorus Boreal coniferous Temperate deciduous Tropical rain forest

0.033–0.060 0.06–0.08 0.2–0.3

0.075–0.15 0.20–0.10 0.001–0.005

0.04–1.06 0.91–1.68 0.06–7.2

300 6 0.6

Potassium Boreal coniferous Temperate deciduous Tropical rain forest

0.15–0.35 0.3–0.6 2.0–3.5

0.3–0.75 0.050–0.15 0.020–0.040

0.07–0.8 0.01–38 0.05–7.1

100 1 0.2


720

in boreal forest soils. Nitrogen limitation of productivity, therefore, is weak in tropical forests and increases from temperate to boreal and tundra forest systems. Trees may obtain organic N and P from the soil via mycorrhizae or by relocation from older foliage prior to abscission, and thereby, partly reduce their dependence on soil as a source of inorganic nutrients. Increased understanding of the fundamental relationships between soil properties and plant nutrients requirements will most likely come from examination of plant–fauna–microbe interactions at root surfaces (rhizosphere), rather than in the bulk soil.

INFLUENCE OF ABIOTIC FACTORS Forests have distinctive physiographic, floristic, and edaphic characteristics that vary predictably across the landscape within a climatically homogeneous region. Differences in the elemental content of parent material influence the tree species composition between and within a landscape unit. For example, wind deposited soils, which support hardwood or mixed wood forest, are likely to be fine textured with high nutrient supplying capacity. In contrast, outwash sands that often support pine forests are coarse textured and infertile. Heterogeneity within the landscape results in sites differing in microclimatic conditions, and physical and chemical properties, which produces different geochemical reaction rates and pools of available nutrients in soil. Soil type and topographic–microclimate interactions are important feedbacks that influence biological processes, such as the rate of N mineralization in soil. Low P availability is a characteristic of geomorphically old, highly weathered tropical, subtropical, and warm temperate soils.[3] The type and age of parent material from which the soil is derived can influence the base status and nutrient levels in soil. Soils in glaciated regions are relatively young and rich in weatherable minerals. Mineral weathering is an important source of most nutrients for plant uptake, with the exception of N. Nutrient availability is regulated by the balance between weathering of soil minerals and precipitation, adsorbtion, and fixation reactions in soil. Edaphic conditions can exert a strong influence on forest productivity and produce considerable variation in nutrient cycling processes. Soils with low N, P, or pH support trees with low litter quality (high in lignin and tannins that bind N) that decomposes slowly. Edaphic limitations on growth may be compensated for by an increase in rooting density and depth. Some late-succession or tolerant species have a shallower root distribution relative to intolerant pioneer species and are adapted to sites where nutrients and moisture



are concentrated at the soil surface. In contrast, nutrient uptake from sub-soil horizons is more important in highly weathered warm temperate soils where nutrient depletion takes place deeper in the soil.

ROLE OF DISTURBANCE Disturbances such as fire, harvesting, hurricanes, or pests affect nutrient cycling long after the event. In fire-dominated ecosystems, intensive wildfire results in a horizontal and vertical redistribution of ecosystem nutrients. Redistribution results from the combined effects of the following processes: 1) oxidation and volatilization of live and decomposing plant material; 2) convection of ash particles in fire generated winds; 3) water erosion of surface soils; and 4) percolation of solutes through and out of the soil. The relative importance of these processes varies with each nutrient and is modified by differences in fire intensity, soil characteristics, topography, and climatic patterns. Expressed as a percentage of the amount present in vegetation and litter before fire, the changes often follow the order: N > K > Mg > Ca > P Harvesting removes nutrients from the site and interrupts nutrient cycling temporarily. The recovery of the nutrient cycle from harvest disturbance is dependent partly on the rate of re-establishment of trees and competing vegetation. Re-vegetation may occur within months in the tropics, 2–5 years in temperate regions, and longer in boreal and tundra regions.[5] Recovery assumes that the soil’s ability to supply nutrients to plant roots has not been altered by disturbance. If nutrients cannot be supplied by the soil at rates sufficient to at least maintain the rate of growth of the previous forest then fertilization may be necessary to maintain site productivity. Nutrient cycling and the impacts of disturbance on nutrient cycling, have been examined thoroughly in many representative world forests. The impact of natural disturbances and management practices on nutrient cycling processes are generally characterized of the stand or occasionally on a watershed basis. There is a growing demand from policy makers and forest managers for spatial estimates on nutrient cycling at local, regional, and national scales. The availability of N, P, and K in soil largely determines the leaf area, photosynthetic rate, and net primary production of forest ecosystems. Forest management practices that produce physical and chemical changes in the soil that accentuate the cycle of nutrients between soil and trees, may increase


forest productivity. Clear-cut harvesting and site preparation practices (mechanical disturbance, slash burning) remove nutrients from soil in tree components and by increased surface runoff, soil erosion, and off-site movement of nutrients in dissolved form or in sediment transport. In the tropics, potential negative impacts associated with complete forest removal and slash burning are greatest because a larger proportion of site nutrients are contained in the living biomass. Environmental impacts associated with clear-cutting and forest management in general, are confounded by climatic, topographic, soil, and vegetation diversity associated with the world’s forests. Best forest management practices can be utilized to control negative impacts on nutrient cycling.


721

REFERENCES 1. Glossary of Soil Science Terms; Soil Science Society of America: Madison, WI, 1996; 134 pp. 2. Van Cleve, K.; Chapin, F.S., III; Dyrness, C.T., III; Viereck, L.A. Element cycling in taiga forests: state factor control. Bioscience 1991, 41, 78–88. 3. Vitousek, P.M.; Stanford, R.L., Jr. Nutrient cycling in moist tropical forest. Ann. Rev. Ecol. Syst. 1986, 17, 137–167. 4. Cole, D.W.; Rapp, M. Elemental cycling in forest ecosystems. In Dynamic Properties of Forest Ecosystems; Reichle, E.D., Ed.; Cambridge University Press: London, 1981; 341–409. 5. Keenan, R.J.; Kimmins, J.P. The ecological effects of clear-cutting. Environ. Rev. 1993, 1, 121–144.

Forest Ecosystems: Soils Associated with Major Ken Van Rees University of Saskatchewan, Saskatoon, Saskatchewan, Canada

INTRODUCTION Soil is defined as the unconsolidated mineral or organic material that occurs at the Earth’s surface and is capable of supporting plant growth. Forest soils are those soils that have developed underneath forest vegetation.[1] Forests cover approximately 29% of the world’s land surface and play a key role in most ecosystems except for the tundra, deserts, some grasslands, and wetlands. Forests and forest soils contain about 60% of the carbon contained in the Earth’s land surface and thus are an integral part in the global carbon cycle.[2] Soil development or genesis beneath forest ecosystems is influenced by a number of different factors. These factors include the type and nature of the forest vegetation, parent material and topography, climate, human or organism influence, and the amount of time that these factors have been influencing soil development.[3] All these factors combined together can result in soils with different physical, chemical, and biological properties that are unique to those conditions and provide the framework for how soils are classified. Generally there are two main classification systems used in the world: the U.S. soil taxonomic[4] and the FAO systems.[5] Soils types will vary around the world, but some generalizations can be made based on the major forest biomes that exist today. The major forest biomes include boreal and coniferous forests, temperate deciduous=mixed forests, scrub and woodlands, temperate rain forest, tropical rain forests, and tropical monsoon=deciduous forests. A generalized vegetation map of the world showing most of these forest biomes is provided in Fig. 1. Typical soil types associated with these forests are summarized in Table 1 and discussed later for each major forest biome.

forests are generally acidic and form on sandy deposits from glaciation and range in thickness from >1 m to very shallow (30 cm

a

Classes as defined in Agriculture Canada Expert Committee on Soil Survey, 1987. The Canadian System of Soil Classification.

under coniferous forests than deciduous forests: dense conifer canopies intercept much of the incoming solar radiation, thereby reducing the solar heating of the soils underneath, especially of wet soils. Delayed soil heating often implies delayed root growth, thereby shortening the effective growing season and hence forest growth. Fine roots are also sensitive to soil acidification:[7] slowly soluble soil Al may convert to toxicologically active Al as the forest soil acidifies either naturally during the course of their development, or on account of acid deposition.

INFLUENCE OF NUTRIENT SUPPLIES AND AVAILABILITIES Matters dealing with nutrient supplies, retention and availability are of particular importance to sustainable forest production, as follows:

Long-term nutrient supplies of forest ecosystems need to be sustainable in principle. Additions of nutrients to each site come from external sources (atmospheric accretions, liming, and fertilizer), and from internal sources (weathering of primary soil minerals, biochemical cycling). Under natural conditions, origin and type of soil parent material are strong determinants of rate of weathering and related nutrient supplies.


Nutrient retention and release involves numerous reactions, including ion sorption and desorption on mineral and organic colloids; precipitation by metal oxides; complexation by humic substances; immobilization of nutrient ions by microorganisms and SOM mineralization. All these processes are influenced by soil texture, SOM content, cation exchange capacity (CEC), pH, base saturation, aeration, and prevailing soil temperature and moisture regimes.

Nutrient availability results from the net cumulative difference between supplies, immobilization, leaching, mineralization, and past uptake, and is also controlled by chemical equilibria. Potential fertility of a soil can provisionally be appraised from simple field observations. For example, favorable contents and distribution of SOM and clay imply improved N, P and base cation supply, and the presence of black minerals in the parent material may indicate K and Mg sufficiency. To be reliable, such observations must be supplemented by appropriate soil analysis. The effect of soil parent material, via the continuing weathering of primary minerals, is often reflected in the species composition of natural forests. For example, in central regions of eastern North America, soils derived from calcareous shales often support, e.g., basswood, white ash, yellow-poplar and hickory, while soils

Forest Soils Properties and Site Productivity

731

Fig. 1 Soils with contrasting productivity in the Atlantic Region of Northeastern North America. Left: brunisolic podsol supporting vigorous forest of northern tolerant hardwoods. Right: gleyed podsol supporting a pure black spruce stand. Note uniformly brown color, structure (granular to subangular blocky), and abundance of roots throughtout.

derived from acid sedimentary rocks support high percentages of beech, yellow birch, red maple and certain oaks. As well, white cedar and eastern red cedar grow better on calcareous soils than on acid soils. In Sweden, site productivity is grouped by calcium content of the soil parent material: soils derived from calcium-poor substrates commonly support Scots pine at low productivity; soils with intermediate calcium availability support pine and mixed conifer forests with high productivity; soils derived from basic igneous sedimentary substrates support productive stands of Norway spruce and hardwoods.

CA, MG, K, P AND TRACE ELEMENTS In areas where soils are very old and deeply weathered (e.g., tropical and subtropical regions in Africa, Asia, Australia, and Central and South America), or where soils have been heavily cropped, forest productivity is often limited because of limited supplies of Ca, Mg, K, P and=or micronutrients such as Zn, Cu, and B.[8] Such soils often consist of abundant accumulation


of Al and Fe sesquioxides, which typically produce soils with high anion exchange capacities (AEC), but low CEC. As a result, nutrient elements such as K, Mg, and Ca are generally in short supply and severe losses of forest productivity are likely to occur when such soils are stripped of their natural forest vegetation. Such soils, furthermore, may loose their originally friable consistency with gradual SOM loss and become cementation when exposed to air and allowed to dry. Areas that have been covered by glaciers in the recent past and currently have a temperate to boreal climate can support productive forests set within the limits set by climate and soil depth. This can be attributed to the continuing release of Ca, Mg, K, and P through natural weathering of the ground-up bedrock (till). Soils in these areas generally have low accumulations of Fe and Al sesquioxides, except in areas of high soil acidity where forest productivity may decrease as a result of advanced podsolization, i.e., the formation of Fe- and organic matter cemented hardpans within the B horizon. Except for N, mineral deficiencies are infrequently encountered in these areas.

732

Forest Soils Properties and Site Productivity

N AVAILABILITY

REFERENCES

In boreal and sub-boreal climates, forest productivity is commonly restricted by low N availability, which is the result of slow SOM turn-over rates. Under these conditions, the forest floor is the principal source of available N: the A horizon is—more often than not—strongly leached, poorly developed or absent, and N in the B layers is essentially unavailable. In these soils, N availability is further adversely affected by high soil acidity, impeded drainage, and restricted soil aeration. In contrast, soils with mesic or warmer temperature regimes have SOM-enriched A layers with high levels of available N. As shown in many studies,[2] site quality on these soils rises with increasing depth of the A layer.

1. Pritchett, L.W.; Fisher, F.R. Properties and Management of Forest Soils, 2nd Ed.; Wiley: New York, 1987; 494 pp. 2. Carmean, W.H. Forest site quality evaluation in the United States. Adv. Agron. 1975, 27, 209–269. 3. Klinka, K.; van der Horst, W.D.; Nuszdorfer, F.C.; Harding, R.G. An ecosystematic approach to forest planning. For. Chron. 1980, 56, 97–103. 4. Zahner, R. Water deficit and growth of trees. In Water Deficits and Plant Growth; Kozlowski, T.T., Ed.; Academic Press: New York, 1968; 191–254. 5. Drew, M.C. Soil aeration and plant root metabolism. Soil Sci. 1992, 154, 259–268. 6. Wilde, S.A.; Iyer, J.G.; Tanzer, C.; Trautman, W.L.; Watterston, K.G. Growth of Wisconsin Coniferous Plantations; Research Bulletin 262, University of Wisconsin: Madison, WI, 1965; 80 pp. 7. Reuss, J.O.; Walthall, P.M.; Roswall, E.C.; Hopper, R.W.E. Aluminum solubility, calcium–aluminum exchange, and pH in acid forest soils. Soil Sci. Soc. Am. J. 1990, 54, 374–380. 8. Stone, E.L. Microelement nutrition of forest trees: a review. In Forest fertilization—Theory and Practice; Tennessee Valley Authority: Knoxville, TN, 1968; 132–175. 9. Burger, J.A.; Kelting, D.L. Using soil quality indicators to assess forest stand management. For. Ecol. Mgmt. 1999, 122, 155–166. 10. Kimmins, J.P.; Scoullar, K.A. FORCYTE 10; Faculty of Forestry, University of British Columbia: Vancouver, BC, 1983; 112 pp. 11. Bhatti, J.S.; Foster, N.W.; Oja, T.; Moayeri, M.H.; Arp, P.A. Modelling potentially sustainable biomass productivity in jack pine forest stands. Can. J. Soil. Sci. 1998, 78, 105–113.

CONCLUSIONS Forest growth restrictions occur on soils with nutrientpoor parent materials, inefficient water and nutrient storage and retrieval capacities, unfavorable pH, low SOM contents, and extreme texture (sand, clay). Some of these soil restrictions can—in part—be corrected by silvicultural means such as change of forest cover type, control of stand density, fertilization, irrigation, and artificial drainage. All of these growth restrictions have been considered in recent efforts to develop soil quality indices for monitoring sustainability of current forest management practices.[9] Additional efforts are being made to quantify the relationships between soil and forest production by way of modeling.[10,11]


Forest Soils: Calcium Depletion Tom G. Huntington United States Geological Survey (USGS), Augusta, Maine, U.S.A.

INTRODUCTION Calcium depletion in forest soils is the process by which Ca stored in the soil in exchangeable, organic, and mineral-bound forms is gradually lost in response to natural and anthropogenic mechanisms. The primary natural mechanism is the weathering of soil minerals and leaching of dissolved Ca to ground and surface water. The primary anthropogenic mechanisms are removal of Ca in forest products and the acceleration of leaching losses by elevated atmospheric acidic deposition.[1,2] When the rate of Ca losses through leaching and harvest removals exceeds the rate of replenishment through atmospheric deposition and the weathering of primary minerals, soil Ca is depleted. These input and output processes are shown schematically in Fig. 1. Calcium depletion results in decreased Ca availability to plants and lower Ca concentrations in soil, surface, and ground water.

WHAT IS THE EVIDENCE FOR CALCIUM DEPLETION? There is a growing body of evidence that supports the hypothesis that forest soils in the eastern United States and in Europe are experiencing Ca depletion. The evidence can be grouped in three major categories: 1) direct re-measurements over time; 2) mass balance studies; and 3) indirect evidence that is consistent with decreasing Ca. Several studies in Europe[3–5] and the US[6–9] have found significant decreases in exchangeable Ca in forest soil following decades of acidic deposition and uptake by aggrading forests. These re-measurement studies are the most direct evidence that forest soils in many regions are experiencing net Ca depletion. Many studies have measured the input and output fluxes of Ca in forest ecosystems and concluded that the rates of loss substantially exceed inputs.[10–13] Examples of pools and fluxes of Ca determined in biogeochemical studies of this kind are summarized in Table 1. In these studies, investigators have attempted to measure atmospheric inputs, vegetation uptake, and soil leaching losses over a period of several years to determine average annual rates. The rate of weathering is usually estimated indirectly. There is Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120018495 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

more uncertainty in the weathering estimate than in other estimates but it can frequently be constrained using silica export of strontium isotope ratios. The balance of evidence from the majority of intensively studied sites indicates that Ca outputs exceed inputs. The potential for weathering resupply remains an important uncertainty, but the weight of evidence suggests that declines in Ca reserves in forest ecosystems are ongoing, arguing that weathering rates are too slow to replenish Ca.[13,14] There have also been a substantial number of studies that have provided insight into the status of forest soil Ca using indirect evidence. Strontium isotope ratio analysis has been used to infer ongoing Ca depletion in the northeastern United States.[15,16] Trends in stem wood chemistry have also been shown to be consistent with Ca depletion in red spruce dominated forests of the northeastern United States.[17,18] Studies have also reported that the failure of stream water to recover from acidification despite large reductions in sulfur deposition is indicative of Ca depletion.[19] Long-term trends in stream water chemistry and sulfate deposition at several forested watersheds at the Coweeta Hydrologic Laboratory in North Carolina and in the Shenandoah National Park in Virginia support the hypothesis that soil retention of atmospherically derived sulfate has decreased in recent years[20,21] supporting the mechanism for ongoing Ca depletion.[1,2] Many streams in the eastern United States have experienced statistically significant decreases in Ca concentrations that may be related to soil depletion.[10] Other studies have demonstrated positive responses in soil and tree health and vigor following fertilization with Ca suggesting that Ca was a limiting nutrient at these forest sites.[22–24]

WHY IS CALCIUM DEPLETION A THREAT TO FOREST ECOSYSTEMS? Calcium depletion has been implicated in the failure of stream water alkalinity to recover in spite of significant decreases in acidic deposition.[19,25,26] Decreases in Ca availability have been suggested as factors influencing health and vigor of fish,[27] snails,[28] birds,[29] and mollusks.[30] Calcium depletion is of concern for a 733

734

Forest Soils: Calcium Depletion

Atmospheric Deposition of Ca Ca

Rain and Dust Ca

Ca2+

Ca2+ Ca2+

Forest Floor

Ca2+

Leaching Loss To Streams

Ca

Ca

Ca2+

Ca2+

Ca2+

Tree Uptake

Ca2+

Ca2+

Ca

Exchangeable Soil Calcium Mineral Soil

Mineral Weathering

Ca

variety of reasons related to the many critical roles that Ca plays in tree physiology and the likelihood that Ca limitation will adversely influence many aspects of forest function.[31] Calcium limitation in forest

Fig. 1 Calcium cycling in forest ecosystems. Inputs to the available pool of exchangeable soil calcium result from atmospheric deposition in precipitation and dust and the weathering of primary minerals. Outputs from this pool result from plant uptake and removal of wood products and leaching from the soil to streams.

ecosystems is thought to adversely influence disease resistance, wound repair, frost hardiness, and lignin synthesis in trees.[31] Calcium depletion is considered to be an especially serious threat to tropical

Table 1 Soil calcium pools, ecosystem fluxes, and net depletion, from selected forest sites in the southeastern United Sates

Predominant species

Soil exchange pool (kg ha1)

Total soil pool (kg ha1)

Net wood incrementb (kg ha1 yr1)

Total atmospheric deposition (kg ha1 yr1)

Soil leaching (kg ha1 yr1)

Net depletion (kg ha1 yr1)

Location

Ref.a

Stewart Co., Georgia

[33]

Pinus taeda L.

840

840

6

3.2

1.1

3.9

Duke Forest, North Carolina

[13]

Pinus taeda L.

2130

4900

14

8.1

8.1

13.6

Panola Mountain, Georgia

[34]

Carya, Quercus, Liriodendron Tulipifera Pinus taeda L.

2200

9500

10

2.3

2.9

10.7

Clemson, South Carolina

[35]

Pinus taeda L.

1450

NDd

3.4

2.9

3.9

4.4

Oak Ridge, Tennessee

[13]

Pinus taeda L.

6900

8600

5.5

5.4

19.2

19.3

Huntington Forest, New York

[13]

Acer saccharum Marsh., Fagus grandifolia, Ehrh., Acer rubrum L., Betula lutea Michx. F., Picea rubens Sarg

606

120,000

2.7

5.1

15

12.6

(Continued)



735

Table 1 Soil calcium pools, ecosystem fluxes, and net depletion, from selected forest sites in the southeastern United Sates (Continued)

Predominant species

Soil exchange pool (kg ha1)

Total soil pool (kg ha1)

Net wood incrementb (kg ha1 yr1)

Total atmospheric deposition (kg ha1 yr1)

Soil leaching (kg ha1 yr1)

Net depletion (kg ha1 yr1)

Location

Ref.a

Thompson Forest, Washington

[13]

Pseudotsuga menziesii, Alnus rubra

1093

46,000

1.5

8.8

13

5.7

Howland, Maine

[13]

Picea rubens Sarg., Abies balsamea

288

59,000

5.1

4.6

8.0

8.5

Sabah, Malaysia

[36]

Acacia mangium

502

727

47e

< 4.0

ND

ND

Nordmoen, Norway

[13]

Picea abies L. Karst

194

29,000

6.1

5.4

9.0

9.8

Turkey lakes, Ontario

[13]

Acer saccharum Marsh., Betula aleghaniensis Britton, Ostrya viginianna Mill. Acer rubrum L.

671

59,000

1.3

9.3

88

80

St.-Hippolyte Quebec

[37]

Acer saccharum Marsh.

ND

ND

ND

3.0

13.3

ND

a

Reference numbers correspond to reference list. Net wood increment assumes that merchantable (stemwood) will be harvested and removed from the site. Whole-tree harvesting would result in greater removals. c GSMNP, Great Smoky Mountains National Park. d ND, not determined. e Estimated from growth at 45 months, 398 kg ha1 was reportedly removed during harvest. (From Ref.[34].) b

forests growing on highly weathered soils in the tropics.[36]

ecosystem recovery following reductions in emission and deposition of acidic compounds.

CONCLUSIONS

ACKNOWLEDGMENTS

Calcium depletion in forest soils is a natural pedogenic process that can be accelerated by harvest removals and acidic deposition. Several lines of evidence support the fact that Ca depletion is an ongoing process in many forest soils. Direct re-measurements have shown Ca depletion in the eastern U.S. and Europe. Mass balance studies strongly suggest outputs exceed inputs in many intensively studied forest catchments. Indirect evidence from strontium isotope analysis, stem wood chemistry, fertilization experiments, and other studies are also consistent with ongoing Ca depletion. Calcium depletion is a threat to forest ecosystems because decreases in Ca in soils can ultimately adversely influence health and vigor of sensitive trees, fish, snails, birds, and mollusks. Ca Depletion is also a threat because it can delay forest and associated aquatic

This analysis was supported by the U.S. Geological Survey’s (USGS) Water Energy and Biogeochemical Budgets, Atmospheric Deposition, and Hydrologic Benchmark Network Programs.


REFERENCES 1. Reuss, J.O.; Johnson, D.W. Deposition and Acidification of Soils and Waters; Springer-Verlag: New York, 1986. 2. Robarge, W.P.; Johnson, D.W. The effects of acidic deposition on forested soils. Adv. Agron. 1992, 47, 1–83. 3. Bergkvist, B.; Folkeson, L. The influence of tree species on acid deposition, proton budgets, and element fluxes in south Swedish forest ecosystems. Ecol. Bull. 1995, 44, 90–99.

736

4. Falkengren-Grerup, U.; Eriksson, H. Changes since soil, vegetation, and forest yield between 1947 and 1988 in beech and oak sites in southern Sweden deciduous forest soils. For. Ecol. Manag. 1990, 38, 37–53. 5. Wesselink, L.G.; Meiwes, K.J.; Matzner, E.; Stein, A. Long term changes in water and soil chemistry in spruce and beech forests, Solling, Germany. Environ. Sci. Technol. 1995, 29, 51–58. 6. Trettin, C.A.; Johnson, D.W.; Todd, D.E.J. Forest nutrient and carbon pools: a 21-year assessment. Soil Sci. Soc. Am. J. 1999, 63, 1436–1448. 7. Knoepp, J.D.; Swank, W.T. Long-term soil chemistry changes in aggrading forest ecosystems. Soil Sci. Soc. Am. J. 1994, 58, 325–331. 8. Richter, D.D.; Markewitz, D.; Wells, C.G.; Allen, H.L.; April, R.; Heine, P.R.; Urrego, B. Soil chemical change during three decades in an old-field loblolly pine (Pinus taeda L.) ecosystem. Ecology 1994, 75, 1463–1473. 9. Johnson, A.H.; Anderson, S.B.; Siccama, T.G. Acid rain and soils of the Adirondacks: I. Changes in pH and available calcium. Can. J. For. Res. 1994, 24, 193–198. 10. Huntington, T.G.; Hooper, R.P.; Johnson, C.E.; Aulenbach, B.T.; Cappellato, R.; Blum, A.E. Calcium depletion in a southeastern United States forest ecosystem. Soil Sci. Soc. Am. J. 2000, 64, 1845–1858. 11. Adams, M.B.; Burger, J.A.; Jenkins, A.B.; Zelazny, L. Impact of harvesting and atmospheric pollution on nutrient depletion of eastern hardwood forests. For. Ecol. Manag. 2000, 138, 301–319. 12. Federer, C.A.; Hornbeck, J.W.; Tritton, L.M.; Martin, W.C.; Pierce, R.S.; Smith, C.T. Long-term depletion of calcium and other nutrients in eastern US forests. Environ. Manag. 1989, 13, 593–601. 13. Johnson, D.W.; Lindberg, S.E. Atmospheric Deposition and Forest Nutrient Cycling; Springer-Verlag: New York, 1992. 14. Turner, R.S.; Cook, R.B.; Van Miegroet, H.; Johnson, D.W.; Elwood, J.W.; Bricker, O.P.; Lindberg, S.E.; Hornberger, G.M. Watershed and Lake Processes Affecting Surface Water Acid–Base Chemistry. In Acid Deposition: State of Science and Technology, Rep. 10, Natl. Acid Precip., Assess. Program, Washington, DC, Nov., 1990; 1990, 1–167. 15. Miller, E.K.; Blum, J.E.; Friedland, A.J. Determination of soil exchangeable-cation loss and weathering rates using Sr isotopes. Nature 1993, 362, 438–441. 16. Bailey, S.W.; Hornbeck, J.W.; Driscoll, C.T.; Gaudett, H.E. Calcium inputs and transport in a base poor forest ecosystem as interpreted by Sr isotopes. Water Resour. Res. 1996, 32, 707–719. 17. Bondietti, E.A.; Momoshima, N.; Shortle, W.C.; Smith, K.T. A historical perspective on divalent cation trends in red spruce stemwood and the hypothetical relationship to acidic deposition. Can. J. For. Res. 1990, 20, 1850–1858. 18. Shortle, W.C.; Smith, K.T.; Minocha, R.; Larwence, G.B.; David, M.B. Acidic deposition, cation mobilization, and biochemical indicators of stress in health red spruce. J. Environ. Qual. 1997, 26, 871–876.



19. Stoddard, J.L.; Jeffries, D.S.; Lu¨kewille, A.; Clair, T.A.; Dillon, P.J.; Driscoll, C.T.; Forsius, M.; Johannessen, M.; Kahl, J.S.; Kellogg, J.H.; Kemp, A.; Mannio, J.; Monteith, D.T.; Murdoch, P.S.; Patrick, S.; Rebsdorf, A.; Skjelkva˚le, B.L.; Stainton, M.P.; Traaen, T.; Van_Dam, H.; Webster, K.E.; Wieting, J.; Wilander, A. Regional trends in aquatic recovery from acidification in North America and Europe. Nature 1999, 401, 575–578. 20. Johnson, D.W.; Swank, W.T.; Vose, J.M. Simulated effects of atmospheric sulfur deposition on nutrient cycling in a mixed deciduous forest. Biogeochemistry 1993, 23, 169–196. 21. Ryan, P.F.; Hornberger, G.M.; Cosby, B.J.; Galloway, J.N.; Webb, J.R.; Rastetter, E.B. Changes in the chemical composition of stream water in two catchments in the Shenandoah National Park, Virginia, in response to atmospheric deposition of sulfur. Water Resour. Res. 1989, 25, 2091–2099. 22. Moore, J.-D.; Camire, C.; Ouimet, R. Effects of liming on the nutrition, vigour, and growth of sugar maple at the Lake Claire Watershed, Quebec, Canada. Can. J. For. Res. 2000, 30, 725–732. 23. Long, R.P.; Horsley, S.B.; Lilja, P.R. Impact of forest liming on growth and crown vigor of sugar maple and associated hardwoods. Can. J. For. Res. 1997, 27, 1560–1573. 24. Nilsson, L.W.; Wiklund, K. Nutrient balance and P, K, Ca, Mg, S and B accumulation in a Norway spruce stand following ammonium sulphate application, fertigation, irrigation drought and N-free-fertilisation. Plant Soil 1995, 168, 437–446. 25. Kirchner, J.W.; Lydersen, E. Base cation depletion and potential long-term acidification of Norwegian catchments. Environ. Sci. Technol. 1995, 29, 1953–1960. 26. Lawrence, G.B.; David, M.B.; Lovett, G.M.; Murdoch, P.S.; Burns, D.A.; Stoddard, J.L.; Baldigo, B.P.; Porter, J.H.; Thompson, A.W. Soil calcium status and the response of stream chemistry to changing acidic deposition rates in the Catskill Mountains of New York. Ecol. Appl. 1999, 9, 1059–1072. 27. Bulger, A.J.; Cosby, B.J.; Webb, J.R. Current, reconstructed past, and projected future status of brook trout (Salvelinus fontinalis) streams in Virginia. Can. J. Fish. Aquat. Sci. 2000, 57, 1515–1523. 28. Graveland, J. Avian eggshell formation in calcium-rich and calcium-poor habitats: importance of snail shells and anthropogenic calcium sources. Can. J. Zool. 1996, 74, 1035–1044. 29. Graveland, J.; van_der_Wal, R.; van_Balen, J.H.; van_Noordwijk, A.J. Poor reproduction in forest passerines from decline of snail abundance on acidified soils. Nature 1994, 368, 446–448. 30. Wareborn, I. Changes in land mollusc fauna and soil chemistry in an inland district in southern Sweden. Ecography 1992, 15, 62–69. 31. McLaughlin, S.B.; Wimmer, R. Tansley review no. 104. Calcium physiology and its role in terrestrial ecosystem processes. New Phytol. 1999, 142, 373–417. 32. Johnson, D.W.; Todd, D.E. Nutrient cycling in forests of Walker Branch Watershed, Tennessee: Roles of


uptake and leaching in causing soil changes. J. Environ. Qual. 1990, 19, 97–104. 33. Huntington, T.G. Assessment of the potential role of atmospheric acidic deposition in the pattern of southern pine beetle infestation in the northwest Coastal Plain Province of Georgia, 1992–1995. In U.S. Geological Survey Water Resources Investigation Report 96-4131; U.S. Geological Survey: Reston, VA, 1996; 75. 34. Huntington, T.G. The potential for calcium depletion in forest ecosystems of southeastern United States: review and analysis. Global Biogeochem. Cycles 2000, 14, 623–638.


737

35. Johnson, D.W.; Kelly, J.M.; Swank, W.T.; Cole, D.W.; Van Miegrot, H.; Hornbeck, J.W.; Pierce, R.S.; Van Lear, D. The effects of leaching and whole tree harvesting on cation budgets of several forests. J. Environ. Qual. 1988, 17, 418–424. 36. Nykvist, N. Tropical soils can suffer from a serious deficiency of calcium after logging. Ambio 2000, 29, 310–313. 37. Hendershot, W.H.; Courchesne, F. Effect of base cation addition on soil chemistry in a sugar maple forest of the lower Laurentians Quebec. Can. J. For. Res. 1994, 24, 609–617.

Future of Soil Science: Role of Soils Richard W. Arnold United States Department of Agriculture, Washington, D.C., U.S.A.

INTRODUCTION Recent studies indicate that during the Anthropocene humanity has experienced exponential growth of consumption, generation of waste, and population.[1] The human economy depends on the planet’s natural capital that provides all ecological services and natural resources, and since about 1980 the human demands have exceeded the capacity of the earth to sustain such use.[2] Cultural attitudes determine the role of soils in today’s world as our fragmented global community struggles to resolve the global issues of food security, environmental protection, and overall sustainability. A variety of world views influences the search for a sustainable, socially acceptable balance among soil functions that provides for viable economic growth and development, safe healthy environments, and intergenerational equity.[3] Linking entire social systems in a web of production, distribution, and consumption, agriculture often foreshadows the degree of economic well-being.[4] Because agriculture operates simultaneously in the realms of ecology and economics, each of which marks time by different clocks, decisions affecting food security and environmental protection have become increasingly complex and variable over time and space.

THE PEDOSPHERE Soils are a critical interface between society and natural resources. Thus, the basic principles of the organization and functioning of the Earth’s soil cover, the pedosphere, can provide a scientific basis for programs addressing global sustainability throughout the 21st century.[5] Natural soils result from the interaction of processes taking place over time on the Earth’s surface. Most involve gases and liquids that transform the solid phase of the surficial materials into features recognized as soils. These processes are influenced by soil-forming factors— namely, climate, biota, topography, parent material, and time—leading to great heterogeneity in the world’s soil cover. Geomorphic processes that alter the landscape by erosion, transport, and deposition of rocks and sediments, and the interaction of biological systems with soils are all subject to major modifications associated Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042624 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

with global and regional climate changes. The result is an intricate patchwork of soils, ranging in size from a few square meters to thousands of square kilometers. For millennia, naturally evolving soils were dominant in the world, and determining the properties and distributions of major kinds of soils and their local geographic associates enabled societies to effectively tap these resources to satisfy their needs. Soils provided habitats for plants, animals, and micro-organisms. Soils possessed the capacity for fertility and potential productivity because of water, physical support, and biological interactions that provided nutrients.[6] Mapping the pedosphere and deciphering the sequence of events and processes causing such complexity have revealed a multidimensional hierarchy that is meaningful for assessing many kinds of soils and predicting soil-related behavior.

ALTERED ECOSYSTEMS For more than a century, scientific studies of soils as natural independent entities on the Earth’s surface have contributed to a better understanding of the interconnectivity of the Earth’s systems.[7] As societies introduced more and more invasive procedures, they drastically altered ecosystem processes and biogeochemical cycles. Although some ecosystems have been enhanced, more have been degraded and are being used unsustainably.[8] Now it has become relevant to monitor, predict, and mediate the behavior and responses of both natural and artificial soil environments and landscapes. It is believed that many changes being made in ecosystems are increasing the likelihood of detrimental nonlinear changes in the future.[1] With increases in the size of human population and its increasing rate of consumption, the available natural resources are stressed, some beyond their limits of resilience. One estimate is that human resource use in 2000 was about 20% above the global carrying capacity.[2] When soils are so stressed, they are unable to return to their former productive states without massive external inputs. Thus, sustainable integration of societal desires and natural resource capabilities is commonly jeopardized. The role of soils, consequently, can be viewed as the set of trade-offs among their various functions, as determined by current society. 741

742

FUNCTIONS OF SOILS Functions of soils commonly include: 1) to promote biomass transformations; 2) to serve as the Earth’s geomembrane to filter and buffer; 3) to provide biological habitats; 4) to provide usable materials; and 5) to serve as sources and repositories of our cultures.[9] Soils are primarily used by most people for the production of food, feed, fuel, and fiber. The capacity to store and release water and the ability to renew, store, and release plant nutrients have dominated agronomic and forestry research and practical experimentation for many years. Biomass transformations are highly dependent on the microbiological populations that inhabit soils and facilitate the formation and use of beneficial compounds. The pedosphere is a sensitive geomembrane, which mediates the transfer of air, water, and energy into, out of, and among the biosphere, atmosphere, hydrosphere, and the lithosphere (Fig. 1). Temperatures are moderated with depth, and moisture and associated compounds are filtered, retained, stored, and transferred in ways that contribute to clean, healthy environments. Soils are the habitat for millions of organisms, ranging from cellular bacteria to burrowing animals. Communities of micro-organisms decompose organic materials, facilitate the release of mineral elements,

Future of Soil Science: Role of Soils

and produce the chemical and biological compounds essential for life on Earth.[6] Many soils are directly used as raw materials for constructing dams and foundations; others are source materials for landscaping industrial and urban sites; and some are ingredients for bricks and ceramic products. Most transportation networks and urban communities rest on soils. In addition, soils are excavated to create disposal sites for society’s numerous wastes. Soils that are disturbed or removed from their natural environments are commonly called ‘‘dirt.’’ As such, displaced soil materials are generally considered a nuisance in our daily activities and need to be washed out, removed, and personal contact with them minimized. Historically, stigma was often attached to those associated with the ‘‘filth and dirt’’ of using and managing soils.[10] In the stories of most indigenous cultures, the Earth is sanctified and revered as a vital element of nature, thereby reinforcing humankind’s link with it. This sacredness remains in the use of soils as cemeteries, and as places where the spirits of ancestors reside. Archeological investigations generally involve deciphering the memories of times and events recorded in soil properties.

CONTINUING NEEDS

Fig. 1 The pedosphere results from the interaction throughout space and time of the atmosphere, biosphere, and hydrosphere with the lithosphere. Once developed, the pedosphere plays a significant role in supporting and maintaining life on the planet.


It is obvious that soils will continue to be used to maintain and improve bioproductivity for food, feed, fuel, and fiber for a long time because the supply of elements and minerals necessary for life are derived, directly or indirectly, from soils. Healthy ecosystems and environments depend on soil-hosted microorganisms that facilitate filtering and purifying. There is still much to be learned about how soils behave in maintaining healthy anthropogenic landscapes. Unprecedented demands to safely handle wastes and provide major increases in food and feed are enormous challenges.[2,11] Perhaps less appreciated is the need of the human psyche for renewal through contact with nature. It appears that psychological well-being is, in part, related to our communion with the beauty, tranquility, and mysterious forces of nature. Although we are connected to the land and soils in ways that most of us do not readily comprehend, we do recognize the sense of belonging and the feeling of renewal associated with our contacts with the natural world. Is it a ‘‘dust thou art, and to dust thou shall return’’ syndrome in which our life cycle is subsumed in other natural life cycles? Whatever the explanation, the human-to-nature relationship is important, and soil is vital to our survival and growth.


CHANGING DEMOGRAPHICS It has been suggested that sometime in this century more than our entire current population will be living in cities.[12] As cities evolve into megacities through the use of adjacent lands, the unique culture of urban dwellers, including their estrangement from rural landscapes and ecosystems, also develops. Ecological functions of soils are common in extensive rural areas of the Earth, where the productive capacity of soils is easily recognized. In urban landscapes, which are very different in appearance, structure, and composition, the role of soils is commonly not observed, or even imagined, except in parks and residential lawns and gardens. Sewers and trash trucks remove numerous waste products, smoke stacks and vehicles exude particulates into the atmosphere, and streams and rivers wash away other debris. In rural landscapes, many of these functions are provided by the pedosphere—the covering of soils that seems to be ubiquitous and so common that it is taken-for-granted. How much relatively undisturbed land would be required to provide the energy and products consumed in an urban environment, and to adequately handle its wastes? Such a measure could be thought of as a city’s environmental footprint.[11] At present we do not measure and monitor the enormous fluxes that occur in urban environments, yet they are becoming major stressors on our global habitat.

INFLUENCING THE ROLE Is there a possibility of simply stretching current ecological theory to encompass urban ecosystems? It can be argued that our understanding of dynamics and processes of populations and soils can be extended to homeowners’ associations and pavement.[11] If people act as other organisms do, guided by individual self-interest, there is no basis for a moral or aesthetic

743

call to environmental stewardship. Is there, perhaps, a spiritual or moral dimension that defies explanations offered by evolution or natural selection? The challenge of understanding urban ecosystems requires specialists from many different disciplines, but it also requires that at least some individuals think in interdisciplinary and multidisciplinary ways—a task that may be difficult to accomplish. The single most important force of landscape change in urban areas is land conversion driven by institutional decisions, population growth, and economic forces.[11] A city’s footprint, which indicates the dependence of an urban ecosystem on other ecosystems, may be tens or hundreds of times larger than the city itself. Soil ecosystems are being altered, and even created, to meet the expectations of urban communities. Consequently, the kinds and patterns of adjacent soil landscapes are important to the monitoring and assessment of environmental health and sustainability.[2,8]

THE FUTURE Humans have made tremendous achievements in their mutual adjustment to accepted standards.[13] For example, at a busy street corner clusters of pedestrians wait, voluntarily, for the light to change to green. When the light changes, surging individuals instinctively negotiate their way across the intersection, without bumping or colliding with those coming from the other direction. The voluntary acceptance of the established standard of traffic regulation by lights results in personal safety. The adaptive flexibility observed in this ‘‘uncentralized’’ operation[13] suggests how the balance of the soil functions will be solved in the future. A global consensus must develop that will clearly define a minimum set of common norms and international standards for sustainable uses of various kinds of soils. Knowledge about the limitations of specific kinds of soils for particular uses is essential for informed decisions and agricultural policies

Table 1 Soil attributes that can become future constraints to achieving desired balances of soil functions if excessive demands are placed on soil resources Soil attribute

Constraint

Resilience

Recovery from disturbance

Productivity

Capability for plant growth and yield

Responsiveness

Capacity for external enhancement

Sustainability

Dynamic equilibrium of interactions

Resistance

Stability to maintain current condition

Flexibility

Multiplicity of uses related to properties

Pedoclimate

Location and extent of suitable climate

Residence time

Capacity to store and release compounds

Geography

Availability due to location or intricacy of pattern


744

(Table 1). For every conceivable use of soil there is a hypothetical ideal soil with the right set of properties and supporting processes to achieve a satisfactory level of success. Comparing local soils with a specified ‘‘ideal’’ soil can lead to recommended measures that minimize limitations and contribute toward attaining the expected behavior of such an ideal soil. When local economics of managing individual soils are considered, it is possible to develop rankings of suitability and economic feasibility. Understanding soil property–soil process relationships can provide a basis for creating and improving soil ecosystems needed to support the constantly changing global environment. International Organization for Standards (ISO), or something similar, will serve as universal ‘‘rules of the road.’’ Individuals and communities will adapt their actions and integration of needs with the available resources. Local flexibility will occur in the mutual adaptations applied to these global standards of ethical and environmental stewardship of the world’s natural resources. All these conditions can be achieved in, and by, the world’s diverse cultures through their own mutual adaptations. The future role of soils is truly in the hands and hearts of the people.

CONCLUSIONS Humanity has been overshooting the resource capacity of our Earth for several decades. There is great disparity among nations in their ecological footprints, thus the challenges of implementing practices to attain sustainability are many. Soils are vital to supplying food, feed, fiber, and fuel to support the development of future generations, consequently improved understanding of functions and limitations of local soils is relevant to helping meet the desires for resources and the handling of wastes in a sustainable global habitat for all humanity.

REFERENCES 1. Meadows, D.H.; Randers, J.; Meadows, D.L. Limits to Growth: The 30-year Update; Chelsea Green Publ. Co.: White River Junction, VT, 2004.



2. Wackernagel, M.; Schulz, N.B.; Deumling, D.; Linares, A.C.; Jenkins, M.; Kapos, V.; Monfreda, C.; Loh, J.; Myers, N.; Norgaard, R.; Randers, J. Tracking the ecological overshoot of the human economy. Proc. Natl. Acad. Sci. (U.S.) 2002, 99 (14), 9266–9271. 3. Karlen, D.L.; Mausbach, M.J.; Doran, J.W.; Cline, R.G.; Harris, R.F.; Schuman, G.E. Soil quality: a concept, definition, and framework for evaluation. Soil Sci. Soc. Am. J. 1997, 61, 4–10. 4. De Soysa, I.; Gleditsch, N.P.; Gibson, M.; Sollenberg, M.; Westing, A.H. To Cultivate Peace: Agriculture in a World of Conflict; PRIO Report 1=99; International Peace Research Institute (PRIO): Oslo, 1999; 1–89. 5. Sokolov, I.A. Spatial–temporal organization of the pedosphere and its evolutionary and ecologic causes. Eurasian Soil Sci. 1994, 26 (4), 10–25. 6. Kovda, V.A. The Role and Functions of the Soil Cover in the Earth’s Biosphere; Scientific Center of Biological Research, Academy of Science, USSR: Pushchino, 1985; 1–12. 7. Targulian, V.O.; Rode, A.A.; Dmitriev, N.A.; Armand, A.D. Soil as a component of natural ecosystems and the study of its history, modern dynamics and anthropogenic changes. In Selection, Management and Utilization of Biosphere Reserves, Proceedings of the United States-Union of Soviet Socialist Republics Symposium on Biosphere Reserves, Moscow, USSR, May 1976; Franklin, J.F., Krugman, S.L., Eds.; USDA Pacific Northwest Forest and Range Experiment Station: Corvallis, Oregon, 1979; General Technical Report PNW 82, 186–197. 8. Millennium Ecosystem Assessment. Main findings; http:== www.greenfacts.org=ecosystems=millennium-assessment3199-main-findings.htm (accessed March 2005). 9. German Advisory Council on Global Change. World in Transition: The Threat To Soils; 1994, Annual Report; Economica Verlag: Bonn, 1995. 10. Yaalon, D.H.; Arnold, R.W. Attitudes toward soils and their societal relevance; then and now. Soil Sci. 2000, 165 (1), 5–12. 11. Collins, J.P.; Kinzig, A.; Grimm, N.B.; Fagan, W.F.; Hope, D.; Wu, J.; Borer, E.T. A new urban ecology. Am. Sci. 2000, 88 (5), 416–425. 12. Brown, L.R.; Gardner, G.; Halweil, B. Beyond Malthus: Sixteen Dimensions of the Population Problems; Worldwatch Paper 143; Worldwatch Institute: Washington, DC, 1998. 13. Cleveland, H. Coming soon: the nobody-in-charge society. Futurist 2000, 34 (5), 52–56.

Gas and Vapor Phase Transport Dennis E. Rolston University of California, Davis, California, U.S.A.

INTRODUCTION Diffusion is the principal mechanism in the interchange of gases between the soil and the atmosphere. The interchange results from concentration gradients established within soil by respiration of microorganisms and plant roots; by production of gases associated with biological reactions such as fermentation and nitrogen transformations; and by soil incorporation of materials such as fumigants, anhydrous ammonia, pesticides, and various volatile organic chemicals in toxic waste sites. The diffusion of water vapor within the soil also occurs due to differences in vapor pressure gradients induced by temperature differences or by evaporative conditions at the soil surface.

TRANSPORT PHYSICS AND EQUATIONS The diffusion velocities of gas mixtures in porous media are related to each other in a complex manner dependent upon the mole fraction of each gas, the molar fluxes of each gas, and the binary diffusion coefficient of each gas pair. If gravity effects are ignored or diffusion occurs only horizontally, the well-known Stefan–Maxwell equations provide the theoretical framework for diffusion of gases in soils. Fick’s law for diffusion is a restrictive case of the Stefan–Maxwell equations and is generally applicable for only a few special cases.[1] One of these cases is for the diffusion of a trace gas in a binary mixture, meaning that the mole fraction of the tracer gas is small. The second special case is for diffusion of two gases in a closed system (total pressure remains constant). In this case, neither gas needs to be in trace amounts. A third case where Fick’s law is applicable is for a three-component system where one of the gases exists in trace amounts and the binary diffusion coefficients of the other two pairs do not differ much from one another (basically the first case). Examples of this case would be for the gas pairs of N2–O2 and N2–Ar. Assuming that the special case conditions are met, Fick’s law is given by Mg dC g ¼ fg ¼ Dp At dx

ð1Þ

Encyclopedia of Soil Science DOI: 10.1081/E-ESS-1-120042692 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

where Mg is the amount of gas diffusing (g gas), A is the cross-sectional area of the soil (m2 soil), t is time (sec), fg is the gas flux density (g gas=m2 soil=sec), Cg is concentration in the gaseous phase (g gas=m3 soil air), x is distance (m soil), and Dp is the soil-gas diffusion coefficient (m3 soil air=m soil=sec).

MEASUREMENT OF THE SOIL-GAS DIFFUSION COEFFICIENT The soil-gas diffusion coefficient, Dp, is the fundamental property that must be known to use Eq. (1) to calculate gas transport in soils. Values of Dp change with soil-air content and tortuosity (crookedness of the diffusion path). The standard laboratory method for measuring the soil-gas diffusion coefficient is based upon establishing gas concentration, Co, within a chamber (Fig. 1). One end of a soil core of concentration Cs is placed in contact with the gas within the chamber.[2] The other end of the soil core is maintained at concentration Cs. The gas of interest diffuses either into or out of the chamber depending upon the concentration Co compared to that outside the core. Obviously, the other gases making up the atmosphere will diffuse in an opposite direction to that of the gas of interest. The time rate of change of concentration in the chamber is related to the soil-gas diffusion coefficient and can be described by equations for unsteady diffusion of gas. Several investigators[2] have used similar procedures. The unsteady diffusion of a gas, which is nonreactive (physically, biologically, and chemically), is described by the combination of Fick’s first law [Eq. (1)] and the continuity equation (conservation of mass)

e

@Cg @ 2 Cg ¼ Dp @t @x2

ð2Þ

where e is the soil-air content (m3 air=m3 soil). In developing Eq. (2), it is assumed that the soil is uniform with respect to the diffusion coefficient and that e is constant in space and time. A simple solution of Eq. (2) that allows for the determination of Dp from laboratory measurements is available.[2,3] 745

746

Gas and Vapor Phase Transport

(m3 voids=m3 soil), and b is the Campbell pore-size distribution parameter, corresponding to the slope of a plot of the log of the soil-water pressure potential vs. the log of the volumetric soil-water content. It has also been shown that there is a significant effect of macroporosity on Dp.[8] In this respect, macroporosity is defined as the air-filled porosity at a soil-water pressure head of –100 cm H2O (–10 kPa), corresponding to the volumetric content of soil pores with an equivalent pore diameter larger than 30 mm. The macroporosity (e100) is found by subtracting the volumetric soil-water content measured at a soil-water pressure head of –100 cm from the soil total porosity. Using this concept results in an additional equation 3 e 2þ3=b Dp ¼ 2e100 þ 0:04e100 D0 e100

Fig. 1 Schematic diagram of the apparatus design used for measuring the soil-gas diffusion coefficient of a soil core.

PREDICTIVE EQUATIONS FOR THE SOIL-GAS DIFFUSION COEFFICIENT Gas transport and fate simulation studies often depend on accurately estimating gas diffusivity (Dp=D0) as a function of soil-air content (e) from easily obtainable physical parameters of the soil, because measured data of Dp(e) are often not available. Many empirical models for predicting the gas diffusivity have been proposed, such as the well-known Millington–Quirk equation,[4] with varying degrees of prediction accuracy.[5] Recently, a series of papers[5–9] have offered new equations that appear to greatly increase the predictive accuracy. Separate equations are needed for undisturbed soils compared to sieved, repacked soils. To accurately predict diffusivity in undisturbed soils, the effects of texture and structure on the diffusivity must be considered.

Undisturbed Soil It has been shown that the Campbell[10] pore-size distribution (water retention) parameter, b, is an effective parameter to describe effects of soil type (soil texture and structure) on Dp(e) in undisturbed soils[5–8] to give e 2þ3=b Dp ¼ F2 D0 F

ð3Þ

where D0 is the gas diffusion coefficient in air (without soil) (m2 air=sec), F is soil total porosity


ð4Þ

To summarize, Eq. (4) appears slightly more accurate and is recommended if both e100 and b are known, whereas Eq. (3) is recommended if only b is known. If the Campbell pore-size distribution parameter, b, is not measured, clay fraction is a good indicator of b,[11,12] and is given by b ¼ 13:6 CF þ 3:5

ð5Þ

Thus, Eq. (5) can be used to determine b to be subsequently used in either Eqs. (3) or (4).

Sieved, Repacked Soil Soil-type effects, such as texture and structure (manifested through pore-size distribution), on gas diffusivity in sieved and repacked soils appear to be minor and can likely be neglected.[9,13,14] A simple, physically based model[9] for Dp(e)=D0 in sieved and repacked soils is given by e Dp ¼ e1:5 D0 F

ð6Þ

The reduction term (e=F) describes the increased tortuosity in a wet soil, compared to a dry soil at the same soil-air content, because of interconnected water films. This equation has been shown to give accurate predictions of the soil-gas diffusion coefficient in sieved, repacked soil samples.[9]

CONCLUSIONS Although techniques for accurately measuring Dp in soil cores are now well established, measured data for Dp are often not available and predictive equations

Gas and Vapor Phase Transport

are required. It is now well established that equations needed to estimate Dp for undisturbed and for disturbed soil samples are indeed different. Equations for undisturbed soil must include soil-physical parameters that take into account the effects of soil type, such as texture, structure, and pore-size distribution, on Dp. Soil-type effects on Dp in sieved and repacked soils appear to be minor.

ARTICLES OF FURTHER INTEREST Aeration Measurement, p. 33. Aeration: Tillage Effects, p. 36. Air Permeability of Soils, p. 60. Greenhouse Gas Fluxes: Measurement, p. 787. Nitrous Oxide Emissions: Agricultural Soils, p. 1129. Nitrous Oxide Emissions: Sources, Sinks, and Strategies, p. 1133. Oxygen Diffusion Rate and Plant Growth, p. 1236. Tillage and Gas Exchange, p. 1773.

REFERENCES 1. Amali, S.; Rolston, D.E. Theoretical investigation of multicomponent volatile organic vapor diffusion: steady-state fluxes. J. Environ. Qual. 1993, 22, 825. 2. Rolston, D.E.; Moldrup, P. Gas diffusivity. In Methods of Soil Analysis, Part 4. Physical Methods; Dane, J.H., Topp, G.C., Eds.; Agronomy Monograph No. 9; Soil Science Society of America: Madison, 2002; 1113. 3. Currie, J.A. Gaseous diffusion in porous media. Part 1. A non-steady state method. Br. J. Appl. Phys. 1960, 11, 314.


747

4. Millington, R.J.; Quirk, J.M. Permeability of porous solids. Trans. Faraday Soc. 1961, 57, 1200. 5. Moldrup, P.; Kruse, C.W.; Rolston, D.E.; Yamaguchi, T. Modeling diffusion and reaction in soils. III. Predicting gas diffusivity from the Campbell soil-water retention model. Soil Sci. 1996, 161, 366. 6. Moldrup, P.; Olesen, T.; Rolston, D.E.; Yamaguchi, T. Modeling diffusion and reaction in soils. VII. Predicting gas and ion diffusivity in undisturbed and sieved soils. Soil Sci. 1997, 162, 632. 7. Moldrup, P.; Olesen, T.; Yamaguchi, T.; Schjønning, P.; Rolston, D.E. Modeling diffusion and reaction in soils. IX. The Buckingham–Burdine–Campbell equation for gas diffusivity in undisturbed soil. Soil Sci. 1999, 164, 542. 8. Moldrup, P.; Olesen, T.; Schjønning, P.; Yamaguchi, T.; Rolston, D.E. Predicting the gas diffusion coefficient in undisturbed soil from soil water characteristics. Soil Sci. Soc. Am. J. 2000, 64, 94. 9. Moldrup, P.; Olesen, T.; Gamst, J.; Schjønning, P.; Yamaguchi, T.; Rolston, D.E. Predicting the gas diffusion coefficient in repacked soil: water-induced linear reduction model. Soil Sci. Soc. Am. J. 2000, 64, 1588. 10. Campbell, G.S. A simple method for determining unsaturated conductivity from moisture retention data. Soil Sci. 1974, 117, 311. 11. Clapp, R.B.; Hornberger, G.M. Empirical equations for some soil hydraulic properties. Water Resour. Res. 1978, 14, 601. 12. Olesen, T.; Moldrup, P.; Henriksen, K.; Petersen, L.W. Modeling diffusion and reaction in soils. IV. New models for predicting ion diffusivity. Soil Sci. 1996, 161, 633. 13. Xu, X.; Nieber, J.L.; Gupta, S.C. Compaction effects on the gas diffusion coefficient in soils. Soil Sci. Soc. Am. J. 1992, 56, 1743. 14. Jin, Y.; Jury, W.A. Characterizing the dependency of gas diffusion coefficient on soil properties. Soil Sci. Soc. Am. J. 1996, 60, 66.

Gelisols James G. Bockheim University of Wisconsin, Madison, Wisconsin, U.S.A.

INTRODUCTION Gelisols, the permafrost-affected soils, comprise 18 million square kilometer or about 13% of the earth’s land surface. They occur in polar regions, the Arctic and Antarctic, and in a few alpine regions (Fig. 1). Gelisols are of global concern because they contain many protected areas and support indigenous people who depend on the land and the surrounding oceans for sustenance. They may be subject to considerable impacts from human development (oil, coal, gas, and gas hydrite exploitation), and are already experiencing global warming.[1] Gelisols have permafrost within 100 cm of the soil surface, or gelic materials within 100 cm of the surface and permafrost within 200 cm of the surface.[2] Permafrost is commonly defined as an earthy material that remains at temperatures below 0 C for two or more years in succession; it may be ice-cemented or, in the case of insufficient interstitial water, dry-frozen. The zone above permafrost that is subject to seasonal thawing is known as the active layer. Gelic materials are defined as seasonally or perennially frozen minerals or organic soil materials that have evidence of cryoturbation (frost churning), ice segregation, or cracking from cryodesiccation.

THE PEDON AS A BASIC SOIL UNIT FOR GELISOLS The pedon is the basic soil unit for sampling in ‘‘Soil Taxonomy,’’[2] and is especially important for describing, classifying, and sampling Gelisols, which often occur in areas with patterned ground. Patterned ground is a general term for any ground surface that is ordered into polygons, nets, circles, or stripes as a result of freezing and thawing processes. Fig. 1 shows earth hummocks in an alpine region resulting from such processes. Each hummock is 1–1.5 m across and is about 1 m high. The pedon is defined so as to encompass the full cycle of patterned ground with a 2 m linear interval or a half cycle with a 2–7 m cycle. This interval is suitable for most patterned ground features such as earth hummocks. In the case of large-scale (>7 m) patterned ground, such as ice-rich, low-center polygons, 748 Copyright © 2006 by Taylor & Francis

two pedons are established: one within the center of the polygon and the other within the trough containing an ice wedge. Therefore, each hummock in the figure represents a single pedon for descriptive and sampling purposes.

CLASSIFICATION OF GELISOLS There are three suborders within the gelisol order, which include histels, turbels, and orthels. They are distinguished on the basis of organic matter content and mineral soils, whether or not there is evidence for cryoturbation. Histels have 80% or more organic materials by volume within the upper 50 cm or to a restricting layer. They are subdivided into five great groups based on the nature of the underlying material and the degree of decomposition: glacistels, folistels, fibristels, hemistels, and sapristels.[3] The glacistel illustrated in Fig. 2 contains 60 cm of organic material directly overlying ground ice. Disturbance to the surface-insulating layer may cause the ice layer to melt and the soil to collapse, a condition known as thermokarst (Fig. 3). The key properties of histels are abundant organic materials ranging from woody to highly decomposed materials, a high moisture holding capacity, and a pH that for an array of histels may range from as low as 2.5 to above 7.0. Although histels comprise less than 3% of the Gelisols globally, they cover large areas in North America and Russia. Turbels represent a second suborder of the gelisols; they are mineral soils subject to cryoturbation. Turbels comprise about 87% of the gelisols on a global basis. Cryoturbation is evidenced by irregular or broken horizons, involutions, organic matter accumulated on the surface of the permafrost, oriented rock fragments, and silt coatings and silt-enriched subsoil horizons that result from freezing and thawing, frost heaving, and cryogenic sorting.[4] Cryoturbation is caused primarily by differential frost heave, but its action can be enhanced by cryostatic pressure, differential swelling, and load casting on poorly drained sites.[5] Fig. 4 shows an aquiturbel on earth hummocks (see also Fig. 1) in northern Canada. This pedon contains an organic layer along the rim of the earth hummocks Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042693 Copyright # 2006 by Taylor & Francis. All rights reserved.

Gelisols

749

Fig. 3 This landform in the Russian far east contains abundant ground ice that has melted following disturbance of the surface organic materials, resulting in a condition known as thermokarst. Fig. 1 Earth hummocks, a type of patterned ground, in the North Cascades of Washington State, U.S.A. Each hummock is 1–1.5 m wide and about 1 m high. This landform features a type of soil known as turbels.

and intensely cryoturbated mineral soil horizons in the center. Although the underlying concept of turbels is that they are subject to cryoturbation, they may contain diagnostic horizons that are common to soils not having permafrost recognized at the great group level. There are seven great groups within the turbels (histoturbels, aquiturbels, anhyturbels, molliturbels, umbriturbels, psammoturbels, and haploturbels) that link them with other soils that do not contain permafrost.[2] It should be emphasized that cryopedogenic processes such as cryoturbation, thermal cracking,

Fig. 2 A hemic glacistel in northern Canada. The organic soil materials, which are intermediate in decomposition from fibric (peat) and sapric (muck), are underlain by ground ice at 60 cm.


and ice segregation are soil-forming processes characteristic of soils with permafrost. They should not be viewed as operating against the other soilforming processes in low-latitude soils; rather, they are distinctive processes producing horizons and properties that are uncommon to other soil orders. Processes common to the other soil orders do operate in gelisols, but at a lesser magnitude, because of the dominance of cryopedogenic processes. The third suborder within the gelisols is the orthels, which include other mineral soils containing permafrost within the upper 100 cm. These soils comprise about 10% of the gelisols globally. The orthels are subdivided into eight great groups that are more or less parallel to those in the turbels and link them to soils not containing permafrost (historthels, aquorthels,

Fig. 4 An aquiturbel developed on an earth hummock in northern Canada (see also Fig. 1).

750

Gelisols

Fig. 6 An apartment complex in the upper Kolyma River region of the Russian Far East that is built on stilts to avoid heat transfer, melting of permafrost, and subsidence.

CONCLUSIONS

Fig. 5 A spodic psammorthel in northern Russia. The soil is excessively drained, developed in sandy floodplain deposits, and contains weakly developed spodic materials in the upper part. (Photo provided by L. Huber.)

anhyorthels, mollorthels, umbrorthels, argiorthels, psammorthels, and haplorthels). Fig. 5 shows a spodic psammorthel in northern Russia. The soil is derived from sandy floodplain deposits and contains weakly developed spodic materials.

GELISOLS AND LAND USE Gelisols offer special challenges to land management for interpretation and practices, including construction (structures and pipelines), mining, forestry, and agriculture. For example, structures must be built on refrigerated pilings or above ground (Fig. 6), so that heat from the structure does not melt the permafrost and cause subsidence. A main concern with gelisols is that they are large C sinks (Table 1). The concern is that warming in the Arctic could increase the thickness of the seasonal thaw layer and enhance heterotrophic respiration, releasing additional CO2 into the atmosphere. In this case, gelisols would become a C source.


Gelisols occur primarily in the Polar Regions and are soils containing permafrost within 100 cm of the soil surface, or gelic materials within 100 cm and permafrost within 200 cm of the surface. Gelisols originate from cryopedogenic processes that include cryoturbation, freezing and thawing, cryogenic sorting, ice segregation, and cryodesiccation. The low temperatures in gelisols give rise to cryostatic pressure development and migration of water with resultant ice buildup. Therefore, gelisols are differentiated from other soils

Table 1 Carbon storage in two gelisols Horizon

Depth (cm)

Organic C (%)

Bulk density C storage (kg/m3) (g/cm3)

Pedon A96-29: typic aquiturbel; Barrow, Alaska[6] Oi 1 25.3 0.42 Bg 32 9.9 1.05 Oejj 2 18.8 0.38 Bgf 15 7.3 0.83 BCgf 12 3.9 1.05 Cgf 38 4.4 1.02

1.1 33.3 1.4 9.1 4.9 17.1

Total

66.9

Pedon 94P0668: typic molliorthel; lower Kolyma River, Russian Far East[2] A 7 7.5 1.04 5.5 AB 5 0.65 1.65 0.5 Bw1 26 0.56 2.4 Bw2 21 0.48 1.7 Bw3 19 0.44 1.66 1.4 Bw4 16 0.44 1.79 1.3 Cf 6 0.39 0.4 Total

13.2

Gelisols

primarily on the basis of thermal characteristics and physical properties that are readily observed in the field. The three main types of gelisols are histels (organic soils with permafrost), turbels (mineral soils with cryoturbation), and orthels (other mineral soils with permafrost in the upper 100 cm). The pedon concept is especially important for sampling gelisols, and is determined by the size of the repeating units of patterned ground features, e.g., 7 m.

751

2.

3. 4.

5.

REFERENCES 1. Bockheim, J.G.; Ping, C.L.; Moore, J.P.; Kimble, J.M. Gelisols: a new proposed order for permafrost-affected soils. Proceedings of the Meeting on the Classification, Correlation, and Management of Permafrost-Affected Soils, Fairbanks, Alaska; Kimble, J.M., Ahrens, R.J.,


6.

Eds.; USDA, Soil Conservation Service, National Soil Survey Center: Lincoln, NE, 1994; 25–44. Soil survey staff. Soil Taxonomy: A Basis System of Soil Classification for Making and Interpreting Soil Surveys, 2nd Ed.; USDA Natural Resources Conservation Service. Agric. Handbook No.436; U.S. Government Printing Office: Washington, DC, 1999. Soil survey staff. Keys to Soil Taxonomy, 9th Ed.; USDA Natural Resources Conservation Service, 2003. Bockheim, J.G.; Tarnocai, C. Recognition of cryoturbation for classifying permafrost-affected soils. Soil Sci. 1998, 162, 927–939. Van Vliet-Lanoe, B. Differential frost heave, load casting and convection: converging mechanisms; a discussion of the origin of cryoturbations. Perm. Periglac. Proc. 1991, 2, 123–139. Bockheim, J.G.; Everett, L.R.; Hinkel, K.M.; Nelson, F.E.; Brown, J. Soil organic carbon storage and distribution in Arctic Tundra, Barrow, Alaska. Soil Sci. Soc. Am. J. 1999, 63, 934–940.

Geology and Soil Ken R. Olson University of Illinois, Urbana, Illinois, U.S.A.

INTRODUCTION Geology is the study of Earth, its internal structure, its materials, its chemical and physical processes, and its physical and biological history.[1] One of the most important geological discoveries was that rocks could form by crystallization of molten material. Rocks do change from one kind to another. As rocks get exposed to the surface weather, particles move downstream, eventually to be deposited as sediments that lithify into sedimentary rocks. By tracing distributions of rock materials, early geologists were able to link sediments to highly consolidated, mineralogically distinct metamorphic rocks.

ATMOSPHERE, HYDROSPHERE, AND LITHOSPHERE The crust, mantle, and core account for 99.97% of the mass of the Earth.[1] The lithosphere includes the Earth’s crust and upper mantle. The term has previously been used to describe the entire portion of the Earth that is composed of rocks. The remaining 0.03% comprises the atmosphere and the hydrosphere. The hydrosphere includes the portion of the Earth’s surface that is covered by water. The hydrosphere behaves as an intermediate reservoir for carbonates, silicates, and other mineral groups leached from the rock of the continents and carried to the oceans by rivers.[1] Both silicates and calcium carbonate follow involved paths from the time they are weathered from continental rock, until they are deposited on the sea floor. The composition of the atmosphere is strongly influenced by the cycling of water from the oceans to the continents, and by its return to the oceans in rivers and subsurface flow. The amount of free oxygen in the atmosphere seems to remain essentially constant. Volcanic gases contain only traces of molecular oxygen but eject large quantities of molecular hydrogen, carbon monoxide, and sulfur dioxide. These gases react with atmospheric oxygen to produce carbon dioxide, water vapor, and sulfur dioxide. The lithosphere, which consumes free oxygen through weathering of rock, and the biosphere, which produces oxygen through photosynthesis, maintain this equilibrium. 752 Copyright © 2006 by Taylor & Francis

The control over this equilibrium may be a feedback mechanism involving the organisms whose by-products and remains become constituents of sedimentary rock.

CYCLING OF ELEMENTS Oxygen, silicon, aluminum, and iron contribute 96% by volume of the elemental composition of the Earth’s crust.[2] The other elements only occupy the remaining 4%. Individual atoms or molecules of such elements such as carbon, nitrogen, oxygen, and sodium change form countless times as they cycle through the atmosphere, biosphere, and lithosphere.[1] Oxygen, for instance, may be converted from a neutral atom to dissolved gas, from combined molecule to ionized particle, to protoplasmic components, and, perhaps, back to the ionic state as a constituent of rock.[1] Other atoms, such as sodium or silicon, may be more constrained in the variety of chemical forms they may assume. These and other representative routes are indicated in a revised and modified schematic shown in Fig. 1.

CHEMICAL WEATHERING PROCESSES ‘‘Weathering’’ is a collective term for the combined effects of all the physical and chemical processes that break down and transform pre-existing rock materials near the Earth’s surface to products that are more stable under the physical and chemical conditions at the surface. Weathering constitutes the response of rock materials to several forms of energy as a function of time.[1] The products of weathering include solids (i.e., sediments and soils) and liquids (including the solutions of salts present in rivers and the ocean). Each physical and chemical factor of weathering affects the outcome of the weathering process. These factors and related variables of the rock cycle, including erosion, transportation, and sedimentation, combine to operate as a complex chemical sorting system which distributes products of different composition to various sites of deposition. Soils are the result of leaching, oxidation, and dissolution of surface materials by the percolation of groundwater and by humic acids derived from oxidized Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042694 Copyright # 2006 by Taylor & Francis. All rights reserved.

Geology and Soil

Fig. 1 Cycling of elements. (From Ref.[1].)

organic materials. New minerals, such as clays, are formed by the chemical alteration of the original mineral of the bedrock.[1] The chemical weathering process develops a soil profile ranging from heavily weathered surface materials down to unaltered rock. The soil zone is the zone of transition between solid rock and the atmosphere. The solid rock, or bedrock, usually has numerous tiny cracks or joints. Chemical weathering caused by water, which fills the cracks, attacks the joints’ surfaces and enlarges the cracks.

PARENT MATERIAL Parent material is the initial or starting material from which a soil develops. This initial material can include either consolidated rock or unconsolidated material deposited by gravity, wind, or water and consists of specific minerals of different sizes, or even plant materials of various plant types.[2] Mineral matter inherited from rocks is referred to as soil parent material because it is the principal ingredient from which soils are formed.[3] However, the principal parent materials of organic soils are decomposing plants. In many cases, relief prior to and during soil formation is related to the nature of the initial soil material.


753

In broad river deltas, crests of natural levees near the stream channels have more coarse material than the areas beyond the levees that are nearly level, and have the finer textured initial material.[4] In steeper topography, where the valleys below the mountain ranges are characterized by broad alluvial–colluvial fans, the initial material near the mountain range contains more coarse and angular material than areas farther away from the mountain range.[5] After unconsolidated parent material is deposited on a stable landscape, or after bedrock has been exposed at the Earth’s surface, soil formation begins and continues over time. The rate of soil formation depends on the climate, including temperature and rainfall. It also depends on vegetation and the activity of other organisms, which live on or in the parent material. These organisms help convert parent materials to soil. Russian pedologists[6] identified parent material as one of the five significant soil-forming factors. Early approaches to soil survey and classification were based on the geology and composition of the soil-forming material.[7,8] The geologic origin and composition of the initial material was identified by the terms ‘‘Agranite soils’’ or ‘‘Aglacial soils.’’ Soils that originate from a parent material inherit the mineral types found in them. Over time, these original minerals are weathered (dissolved) and new minerals form and accumulate in the soils.[2] Russian soil scientists[9] showed the controlling effects of parent material on soil properties. Jenny[10] conducted a systematic analysis of the relationship between soil properties and parent materials from which the soils developed. Jenny proposed parent materials as an independent soil-forming factor, defining parent material as ‘‘the state of the soil system at time zero of soil formation.’’ The physical body of soil and its associated mineralogical and chemical properties are the starting point for the interaction of other soil-forming factors. Previous weathered rock—even a previous soil—could be considered as parent material. The properties of modern soil are the result of the composition of the surficial layer, which existed when the other factors started to impact, and the alterations resulting from the effect of these factors over time. Properties of younger soils are greatly influenced by parent material. Weathering and pedogenic processes result in characteristics of the original parent material being eliminated. Extremely resistant initial material, such as quartz or sand, may still exist in old, weathered soils. It can be difficult to separate the nature (or characteristics) of the initial material and its influence on soil, the kind of ‘‘preweathering’’ of the initial material before becoming parent material for the soil, and the effects of the other soil-forming factors on the parent

754

material of this soil. The complex environmental history (climatic and vegetation changes in the recent geologic past) makes it difficult to separate parent material influences on soil properties from other factor effects. Rock types influence the soil properties of the modern soil. The impact of rock type on soil properties was organized by Buol, Hole, and Mc Craken[11] into the following subdivisions: sedimentary, igneous, and combinations of mineralogically similar metamorphic and igneous rocks. Sedimentary rocks include unconsolidated glacial deposits, loessial deposits, unconsolidated coastal plain sediments, and consolidated rock, such as limestone and dolomites, sandstones, and shales. Siliceous crystalline rocks include more ‘‘acidic’’ quartzose, igneous, and metamorphic rocks including granites, granite gneiss, and schists. Dark-colored ferromagnesian rocks include andosites, diorites, basalt, and hornblende gneiss. Volcanic ash parent materials are composed of noncrystalline materials, any glass fragments, bits of the easily weatherable feldspars, ferromagnesian minerals, and varying amounts of quartz. Mineral components of many soils are inherited almost exclusively from parent materials, while others are developed mostly in situ during the course of weathering and pedogenesis. Primary minerals are formed at high temperature in igneous and metamorphic rocks. Secondary minerals are formed at lower temperature in sedimentary rocks and soils.[12]

CONCLUSIONS The basic model of soil implies that soils are dynamic and geographical. In soil systems, the processes or driving forces are best described as dynamic, rather than static. Morphological properties of soil are the result of processes acting on parent materials. In addressing the influence of parent materials in soil genesis, Chesworth[13,14] stated that ‘‘time has the result of modifying


Geology and Soil

the parent material effects so that only in young or relatively immature soils will the parent material exert its strongest influence on the soil-forming process. That influence will be an inverse function of time.’’

REFERENCES 1. Jackson, C. Geology Today; CRM Books, Communications Research Machines: Del Mar, CA, 1973. 2. Singer, M.J.; Munns, D.N. Soils, An Introduction, 3rd Ed.; Prentice Hall: Upper Saddle River, NJ, 1996. 3. Troeh, F.R.; Thompson, L.M. Soils and Soil Fertility, 5th Ed.; Oxford University Press: New York, 1993. 4. Russell, R.J. River Plains and Sea Coasts; University of California Press: Berkeley, 1967. 5. Birot, P. The Cycle of Erosion in Different Climates; Translated by Jackson, C.I., Clayton, K.M., Eds.; University of California Press: Berkeley, 1960. 6. Dokuchaev, V.V. Russian Chernozem (Russkii Chernozen). In Collected Writings (Sochineniya); Academy of Science: Moscow, USSR, 1883; Vol. 3. 7. von Richthofen, F.F. Fuhrer fur Forschungsreisende; (cited in Joffe, J.S., 1936: Pedology, Rutgers University press, New Brunswick, N.J), 1886. 8. Thaer, A.D. Grundsatze der Rationaellen Landwirtschaft (cited in Joffe, J.S., 1949: Pedology Publ., New Brunswick, NJ), 1809, 1810, 1812; Vol. 1–4. 9. Polynov, B.B. Das Muttergestein als Faktor de Bodenbildung und als Kriterium fur die Bodenklassification. Soil Res. 1930, 2, 165–180. 10. Jenny, H. Factors of Soil Formation; McGraw-Hill: New York, 1941. 11. Buol, S.W.; Hole, F.D.; Mc Cracken, R.J. Soil Genesis and Classification, 2nd Ed.; The Iowa State University Press: Ames, 1980. 12. Jackson, M.L. Chemical composition of soils. In Chemistry of Soil, 2nd Ed.; Bear, F.E., Ed.; Reinhold: New York, 1964. 13. Chesworth, W. The parent rock effect in the genesis of soil. Geoderma 1973, 10, 215–225. 14. Chesworth, W. Conceptual models in pedogenesis: a rejoinder. Geoderma 1976, 16, 257–260.

Geophysics in Soil Science Jeffrey J. Daniels Department of Geological Sciences, The Ohio State University, Columbus, Ohio, U.S.A.

Barry Allred Fdag&Bioen, United States Department of Agriculture (USDA), Columbus, Ohio, U.S.A.

Mary Collins Soil and Water Science Department, University of Florida, Gainesville, Florida, U.S.A.

James Doolittle United States Department of Agriculture (USDA), Newtown Square, Pennsylvania, U.S.A.

INTRODUCTION Geophysics can be defined as the science of measuring physical property changes in the subsurface through the use of instruments located in boreholes, on (or near) the surface, or in the air. The instruments that make these measurements are designed to detect a particular phenomenon caused by a contrast in physical properties. In some cases, there is a source of energy that stimulates a detectable physical change in the subsurface that can be detected by some type of detector that is separate from the source of energy. These methods are called active measurements. Other measurements simply detect a change in the natural background, and these methods are called passive measurements. Geophysical methods can be further classified based on the general category of physical properties that are measured. The general classification includes electrical, seismic, nuclear, potential field, and thermal methods. Detailed summaries of the operational principles of these methods are given in a number of texts in Refs.[1,2].

GEOPHYSICS APPLIED TO SOIL SCIENCE The application of geophysical methods for agricultural purposes has been steadily growing over the past two decades. The impetus for these developments has been the primary need to improve the efficiency of agricultural processes. These improvements have required an increase in the knowledge of the physical and chemical properties at a given field site. The physical properties of interest include the soil texture, moisture, and density, and the location of tiles and drains in the subsurface. The chemical properties of interest include soluble salts, nutrients, and cationexchange capacity of soils. Electrical and neutron Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006627 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

scattering were proven to be the most effective methods in addressing these problems. There are numerous electrical methods, with each method measuring a specific electrical parameter that changes as a function of distribution of electrical permittivity (e), conductivity (s), or magnetic per meability (m) in the subsurface. A summary of the geophysical methods that have been used for the purpose of analyzing the physical and chemical properties of soil are presented in Table 1. The last column lists the soil property that is most commonly the objective of the method. However, other physical and chemical properties will influence the measurement. Electrical Resistivity (Conductivity) The electrical resistivity response (which is the inverse of electrical conductivity) that is measured in soil is affected by a number of different soil characteristics including particle-size distribution, clay mineralogy, cation-exchange capacity, salinity, plant nutrients, moisture content, etc. Therefore soil electrical conductivity mapping can provide indications of spatial trends in these soil characteristics. In the traditional electrical resistivity method, an electrical current is supplied between two electrodes staked into the ground while voltage is concurrently measured between one or more separate pairs of staked electrodes. The method is slow and tedious. Newer technologies with pulled electrode arrays[3] capable of continuous electrical resistivity measurement have made it easier to measure resistivity in the field. One field version of a soil electrical conductivity system was developed by Veris Technologies. This soil electrical conductivity mapping system utilizes direct-contact pulled coulter-electrode resistivity arrays. The OhmMapper TR1 is another resistivity system that is designed for rapid data acquisition. It is a capacitively 755

756

Geophysics in Soil Science

Table 1 Summary of geophysical methods applied to soil analysis problems Method Direct current resistivity

General description Measures variations in subsurface conductivity by measuring the resistance to the flow of direct current.

Frequency range (electrical only)

Soil properties primarily analyzed

0 Hz

Texture

Moisture Induction electromagnetics (EM)

Measures variations in subsurface conductivity by measuring the secondary electromagnetic fields that are induced by low-frequency electromagnetic signal. Induced magnetic and electric fields are both measured.

100 Hz–100 kHz

Texture

Moisture Ground-penetrating radar (GPR)

Measures the propagation (reflection, diffraction, refraction) of a high-frequency electromagnetic wave as a function of travel time.

10 MHz–5 GHz

Texture

Moisture Density Conductivity probe

Measures the electrical conductivity with a probe inserted into the ground.

Moisture Texture

Time domain reflectometry (TDR)

Measures the velocity of a high-frequency electromagnetic wave.

10–100 MHz

Moisture

Neutron–neutron (NS)

Measures the thermal absorption and scattering of thermal neutrons

NA

Moisture Density

coupled resistivity measurement system that operates as a towed dipole–dipole electrode array with continuous data collection. The array is connected to a data logger console and can be integrated with Global Positioning Systems (GPS) while being towed by a person or vehicle. Changing the separation distance between the two dipoles within the array alters the depth of measurement. This capability allows the versatility to produce horizontal maps or vertical profiles of the interpreted electrical conductivity (or resistivity). Photographs of the Vertis and OhmMapper systems are shown in Fig. 1, along with a cross-section profile line of data from the OhmMapper system.

Electromagnetic Induction Electromagnetic induction typically employs an instrument referred to as a ground conductivity meter (GCM). An alternating electrical current is passed through one of two small electric wire coils spaced a set distance apart and housed within the GCM. One coil generates an electromagnetic (EM) field above the surface, inducing a secondary electromagnetic field that propagates into the ground. The second wire coil acts as a receiver measuring the amplitude and phase


components of the fields after they propagate through the ground and is used to calculate an ‘‘apparent’’ value for soil electrical conductivity. Several factors influence the measured response including the primary (instrument) field frequency, and the soil moisture conditions, clay content, and soluble salt contents. The primary advantage of induction methods is that direct contact with the ground is unnecessary. Therefore the method can be used to rapidly acquire measurements of the near-surface conductivity. Electrical Conductivity Probes A variety of hand-operated or machine-powered push probes are available for obtaining point values of electrical conductivity at different depths within the soil profile. Often, these probes contain more than one sensor so that other soil characteristics, such as temperature and penetration resistance, can likewise be measured. Ground-Penetrating Radar Ground-penetrating radar (commonly called GPR) is a high-resolution electromagnetic technique that is primarily designed to investigate the shallow subsurface


757

Fig. 1 Veris 3100 Soil EC Mapping System pulled resistivity array, Ohm apper TR1 vertical resistivity (O m) system, and an interpreted apparent resistivity profile from a test plot located on the Ohio State University Waterman Agricultural and Natural Resources Laboratory in Columbus, Ohio.

of the earth, building materials, and roads and bridges. It is a time-dependent geophysical technique that can provide a 3-D pseudo image of the subsurface, including the fourth dimension of color, and can also provide accurate depth estimates for many common subsurface features. This technology was proven capable of delineating soil profile layers[4] and also shows great promise with regard to updating soil survey information.[5] The radar was used to document the type and variability of soils that occurred in a soil map unit.[6] Collins[7] and Doolittle and Collins[8] demonstrated how soil information can be used as a guide to determine applicability of GPR at a field site and the criteria used to define and classify soils from radar imagery. Today’s GPR technology allows us to model the subsoil in 3-D and geo-reference the data with GPS. Tischler, Collins, and Grunwald[9] demonstrated how GPR data can be incorporated into ArcView and ArcGIS software to create models with GPR and GPS data that show soil subsurface features such as sandy horizons over an argillic (high in clay) horizon.

The details of field data measurement procedures vary from site to site; however, most GPR field measurements are initially displayed in the form of a two-dimensional cross section that is similar to a seismic section (Fig. 2). If the velocity of the electromagnetic wave is known, then a GPR cross section in Fig. 2 can be interpreted as a depth–distance cross section of the subsurface. The two-dimensional cross sections (Fig. 2) provide good information concerning the stratigraphy at a site and the location of the water table, while a threedimensional display can be constructed of closely spaced two-dimensional lines.[10] An example of a GPR amplitude map from a depth interval of 0.7 to 1.2 m (2.3–4 ft) is shown in Fig. 3. The lighter shades on the grayscale map represent locations where a greater amount of reflected radar energy returned to the surface. Locations showing greater reflected radar energy represent dielectric constant discontinuities in the subsurface and often indicate the existence of buried objects. Where there are mapped linear trends

Fig. 2 Examples of applications of GPR field measurements.


758

Fig. 3 GPR amplitude map for depth interval of 0.7–1.2 m showing subsurface drainage system.


environment are associated with water. Neutron probes used for soil moisture measurement emit highenergy neutrons (5 MeV) from a radioactive source, such as radium–beryllium or americium–beryllium. The neutron probes used for agricultural purposes are lowered through the soil profile within boreholes lined with aluminum, steel, or plastic access tubing. When the probe is lowered to measure soil moisture content at a particular depth, the constant emission of high-energy neutrons quickly produces an equilibrium cloud of thermal neutrons surrounding the borehole at that location. The density of this cloud of thermal neutrons depends on the amount of water present in the soil at this depth. The greater the density of the thermal neutron cloud, the greater the amount of water present. Therefore sensors on the probe capable of counting thermal neutrons provide information on soil moisture content at different depths beneath the surface.

of high radar amplitude (energy), subsurface drainage pipes may be present, as shown on this map. CONCLUSIONS Time Domain Reflectometry Time domain reflectometry (TDR) is a high-frequency electromagnetic analysis technique described by Topp, Davis, and Annan[11] The TDR method is based on the fact that the relative electric permittivity (e) of any material is related to the velocity of a high-frequency electromagnetic signal by the relationship: V0 Vm ¼ pffiffi e where V0 and Vm are the velocities of the electromagnetic wave through air and the material, respectively. The relative electric permittivity is equal to the ratio of electric permittivity of the material to the electric permittivity of air (free space). The relative electric permittivity values range from 1 to 81, with a value of 1 being the relative permittivity of air, and 81 being the relative permittivity of water. Hence water is the single most important factor that affects the permittivity of soil. Therefore a measure of velocity of a highfrequency electromagnetic wave can be indirectly used to measure the moisture of soil. Hand-held TDR probes that take electromagnetic pulse travel time measurements and then automatically calculate soil volumetric water content are now readily available. Neutron–Neutron Hydrogen nuclei have a strong capacity for scattering and slowing neutrons. This characteristic can be utilized to measure volumetric moisture content because most of the hydrogen atoms found in the soil


Geophysical methods are important noninvasive tools for determining the physical properties of soil. Electrical and neutron–neutron (NS) methods were demonstrated to be useful in determining soil moisture and in classifying soils. Ground-penetrating radar is a technique that can be used to map soil horizons, bedrock, stratigraphic layers and cultural features (e.g., field tile) in the subsurface. Geophysical methods can be used to rapidly evaluate the conditions of the subsurface over large areas. Geophysical methods have been applied to problems in determining the moisture of soil in the near surface by inserting probes into the ground. The primary methods include the following: 1) neutron–neutron (sometimes called neutron scattering; or NS); 2) time domain reflectometry (TDR); and 3) capacitance measurements. Each of these measurements employs a different physical principle to estimate moisture. All geophysical methods must be calibrated for local conditions, and assume a fairly uniform distribution of the mineralogy throughout the volume that is being analyzed.

REFERENCES 1. Reynolds, J.M. An introduction to applied and environmental geophysics. In Geophysical Methods; Sheriff, R.E., Ed.; Prentice Hall, Wiley: New York, 1989. 2. Sheriff, R.E. Geophysical Methods; Prentice Hall: New York, 1989; 605 pp. 3. Sorensen, K. Pulled array continuous electrical profiling. First Break 1996, 14 (3), 85–90. 4. Kung, K.-J.S.; Boil, J.; Selker, J.S.; Ritter, W.F.; Steenhuis, T.S. Use of ground penetrating radar to


improve water quality monitoring in the vadose zone. In Preferential Flow; Gish, T.J., Shirmohammadi, A., Eds.; ASAE: St. Joseph, MI, 1991; 142–149. 5. Schellentrager, G.W.; Calhoun, T.E.; Doolittle, J.A.; Wettstein, C.A. Using ground penetrating radar to update soil survey information. Soil Sci. Soc. Am. J. 1988, 52 (3), 746–752. 6. Doolittle, J.A.; Schellentrager, G.W. Soil Survey of Orange County, Florida. USDA-SCS; U.S. Government Printing Office: Washington, DC, 1989. 7. Collins, M.E. Soil taxonomy: a useful guide for the application of ground-penetrating radar. 4th International Conference on Ground Penetrating Radar, 1992; Geological Survey of Finland, 1992; 125–132; Special Paper 16.


759

8. Doolittle, J.A.; Collins, M.E. Use of soil information to determine application of ground penetrating radar. J. Appl. Geophys. 1995, 33, 101–108. 9. Tischler, M.A.; Collins, M.E.; Grunwald, S. Integration of ground-penetrating radar data, global positioning systems, and geographic information systems to create three-dimensional soil models. Proc. Ninth Int. Conf. on GPR; Santa Barbara, CA, 2002; 313–316. 10. Daniels, J.J.; Grumman, D.; Vendl, M. Coincident antenna three dimensional GPR. J. Environ. Eng. Geophys. 1997, 2 (1), 1–9. 11. Topp, G.C.; Davis, J.L.; Annan, A.P. Electromagnetic determination of soil water content: measurements in coaxial transmission lines. Water Resour. Res. 1980, 16 (3), 574–582.

Global Resources Paul Reich Hari Eswaran United States Department of Agriculture-Natural Resources Conservation Service (USDA-NRCS), Washington, D.C., U.S.A.

INTRODUCTION A biome is defined as ‘‘a community of organisms interacting with one another and with the chemical and physical factors making up their environments.’’[1] For most purposes, the term biome is used to identify the natural habitat conditions around the world. Depending on the purpose, the global ecosystem is divided into units each characterized by a specific combination of climatic factors. Two major determinants of biome type are precipitation (total and its distribution) and air temperature. These two elements of climate have been commonly used to define the major biomes of the world. A third variable that affects the habitat type is the soil. This section presents a general overview of the soils characterizing the major biomes of the world. Detailed maps showing the biomes of the world are not available due to conceptual differences of definitions and reliable global databases. There are many excellent and detailed studies of specific habitats around the world, and using these and the global soils and climate database of the world,[2] a map showing the distribution of the major biomes was drawn. The terms used to describe the biomes are common in use, but their subdivisions are based on important differentiating factors. MAJOR BIOMES The five major biomes of the world and their subdivisions are listed in Table 1 and their distribution is shown in Fig. 1 Some geographers have used elevation as a differentiating factor and recognized a ‘‘montane biome.’’ This biome’s small extent precludes its description in this section. Few geographers recognize the Mediterranean biome as presented here as a subdivision of the Temperate biome. Within each biome, some major distinction is made. In the Polar biome, an attempt is made to differentiate areas with permafrost and warmer areas with intermittent permafrost. The latter is termed ‘‘interfrost.’’ The Boreal Forest biome has a humid and a semiarid counterpart similar to the temperate and tropical Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042695 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

zones. The semiarid vegetation is generally grass or shrubs. The desert areas are divided into hot, cool, and cold equivalents. Deserts occupy about 30% of the global ice-free landmass. Potential evapotranspiration exceeds precipitation during most days of the year, and the biome shows the greatest contrast in temperature, both diurnal and seasonal. Desert vegetation is very sparse with a few isolated shrubs dominated by succulents and annuals. Grasses and forbs become more dominant at the margins. The tropics occupy about 27% of the landmass and are characterized by only a small variation in temperature during the year. The soil moisture conditions range from semiarid to humid. The availability of soil moisture determines the biota. The temperate grassland and forest biome occupies about 15%. The cold biomes, the boreal and polar, are mainly in the northern latitudes with small areas in South America. The low temperatures control the flora and fauna, and vegetation is scarce to nonexistent in the extreme cold polar regions. Each of these areas has specific kinds of soils that have developed as a response to the specific soil moisture and temperature conditions. The diversity in vegetation has its equivalent in fauna with animal species adapting to the specific bio-climatic conditions. Within each of the biomes, there are specific conditions that promote unique flora and fauna. Examples of such localized systems are volcanic hot springs, wetlands, and oases. Soils of the Major Biomes Soil temperature and moisture, with their seasonal and annual variations, are an integral part of the soil classification system called Soil Taxonomy[3] which is used here. As Fig. 1 and Table 1 are based on soil moisture and temperature conditions, there is an implied link between the biomes and their soil resource endowments. The purpose of this introductory section is to present the general geographic distribution of the biomes and the major soils characterizing each. The major soil orders are listed in Table 2. The subsequent sections elaborate on the soil resources of each of the six biomes listed in Table 1. 765

766

Global Resources

Table 1 Major global biomes with subdivisions and their areas Global land area Biome

Subdivision

Remarks

In thousand km2 10,547 9,685

8.06 7.40

15.46

3,552 9,446

2.72 7.22

9.94

In%

In%

Polar

Permafrost Interfrost

Mosses and lichens Shrubs and stunted trees

Boreal

Semiarid Humid

Shrub Forest

Temperate

Semiarid Humid Mediterranean warm Mediterranean cold

Grassland Forest Shrubs and forest Forbs and shrubs

7,348 12,453 3,624 791

5.62 9.52 2.77 0.60

15.14

Desert

Hot Cool Cold

Barren Barren Barren

4,418 28,521 5,599

3.38 21.81 4.28

29.47

Tropical

Semiarid Humid

Grassland, savanna Forests

20,299 14,514

15.52 11.10

26.62

130,796

100.00

Total

Polar Biome Characterized largely by a mean annual soil temperature of less than 0 C, the Polar biome occupies a large land mass adjacent to the Arctic circle to about 60 N latitude. This is generally the ice-free land of the northern latitudes. The sub-zero soil temperatures that prevail during most of the year are conducive to the formation of

permafrost. The dominant soils of the region are Gelisols, occupying 59% of the Polar biome. Turbels freeze and thaw once or more during a year which triggers cryoturbation or physical mixing of the soil material. The Orthels are generally shallower or drier soils and have little or no cryoturbation. The low temperatures and periodic moisture saturation promote the accumulation of organic matter. The Histels are characterized by the

Fig. 1 Major biomes.


3.37

Global Resources

Table 2 Dominant soils in the major biomes (Area in thousand km2) Polar

Boreal

Temperate

Desert

Tropics

Ice-free land

Permafrost

Interfrost

Semiarid

Humid

Semiarid

Humid

Mediterranean warm

Mediterranean cold

Warm

Cool

Cold

Semiarid

Humid

Gelisols

11,869

10,476

1,393

—

—

—

—

—

—

—

—

—

—

—

Histosols

1,526

—

—

299

686

17

8

—

Spodosols

4,596

55

1,058

227

2,566

105

450

39

6

—

975

6

56

47

196

26

135

32

—

Soil Order

Andisols Oxisols

9,811

—

—

Vertisols

3,160

—

—

2

—

60

76

—

1

14

Aridisols

15,629

Ultisols

11,053

Mollisols Alfisols

1 —

9,161 12,621

9

—

— 15

82

1

—

17

23 —

28

10 651

0

353

18

203

23

1,824

11,014

2,013

266

3,122

20

18

—

—

—

1,110

2,255

136

176

—

—

4,165

638

—

3,566

2,502

342

3,225

1,093

1,059

511

1,280

1,232

874

323

1,002

1,816

1,867

2,149

860

125

49 —

—

6,751

167

3,333

1,636

3,078

683

89

—

—

—

280

172

232

1,173

1,697

807

45

1,708

10,693

— 9,685

87

4

239

147

—

6,199

1,170

27

0

—

10,547

3,387

—

99

21,805 7,122

235

184

21,467 130,796

61

198

303

—

—

—

281

9

Entisols Total

14

37

595

Inceptisols Miscellaneous

93

516

1

20

4

6

153

553

5,016

853

3,552

9,446

7,348

12,453

3,624

791

4,418

28,521

5,599

44

1

4,371

3,240

—

—

20,299

14,514

767


768

high organic matter and occupy about 5% of the Polar biome. These organic soils form the largest contiguous extent of such soils in the world and serve as an important sink of CO2. Other soils that occur in the Tundra zone are Inceptisols 33%, Spodosols 6%, while Mollisols and Entisols are each 1% of the area.[4] Boreal Forest Biome Bordering the southern flank of the Polar biome is the Boreal Forest biome in the northern hemisphere. The high volcanic area between Chile and Argentina has similar climatic regimes but the nature of the soil and physiography may result in a different set of habitat conditions. There are about 243,000 km2 of Andisols or volcanic ash soils characterizing the Boreal Forest biome in the southern hemisphere. Such soils only occur as a small area in the Kamchatca Peninsula in the Northern Hemisphere. In Northern Europe and Siberia, the Boreal Forest biome has a humid part and is bordered in its southern periphery by a semiarid to arid part. In Canada, the semiarid part is in the middle of the continent. A range of soils are present and the most extensive are the Spodosols, which occupy 2.7 million km2 (20.6%). The Spodosols are mostly in the humid part of the Boreal Forest biome and form under acid vegetation. Histosols are present in the depressions in this cold region. Temperate Grassland and Forest Biome The Temperate Grassland and Forest biome extends from about 25 to 55 N with counterparts in the Southern Hemisphere. Large areas in this belt are also deserts. Due to the favorable climate and soil endowments of this biome, much of the land is used for agriculture and native habitats are local and sporadic. About 50% of the zone (12.1 million km2) is occupied by Mollisols and Alfisols (grasslands), and Ultisols in the forests. Their general good fertility and tilth have made them in great demand for grain production. Due to the long history of civilization in the Temperate Grassland and Forest biome of Europe and China, pristine ecosystems are rare. In large areas, there have been successive replacements and changes in the floral and faunal composition. Within the Temperate biome, are areas characterized by moist winters and dry summers. These Mediterranean conditions are conducive to unique ecosystems. These areas occur around the Mediterranean Sea and small areas in the western U.S. and southern Australia. Desert Biome Deserts occupy about 38.6 million km2 and may be distinguished as warm (or tropical), cool (or temperate)


Global Resources

and cold (or boreal) kinds of habitats. Though the biome is characterized by lack of moisture for normal vegetative growth of most plants, the ambient temperature conditions further distinguish the habitat conditions. About 38.6% of the Desert biome is occupied by Aridisols, which by definition have some kind of subsurface horizon. Shifting sands or moving dunes occupy about 13.7% and Entisols, which are very recent deposits, occupying about 33% of the Desert biome. The harshness of the environment has resulted in plants and animals with special adaptive features. Many of the Aridisols are increasingly used for agriculture when irrigation facilities are made available. In this fragile ecosystem, both soil and habitat conditions are drastically altered when irrigation is introduced. Tropics Biome An absence of winter and summer temperature extremes characterizes the Tropics biome, which occupy about 34.8 million km2. Availability of moisture separates the semiarid from the humid tropics. As a corollary to the desert, in the humid tropics there are the perhumid areas where the potential evapotranspiration never exceeds the precipitation during any month of the year. This is a biomic condition that deserves much greater detailed studies. The Tropics biome is the home to the Oxisols (9.6 million km2), which are unique to this biome. These are highly weathered soils where the original vegetation is closed or open forests. Most of the plant nutrients are concentrated in the top 5 cm of the soil and recycled through plant uptake and leaf fall. Ultisols also occur in this biome. Apart from these soils, there are small areas of most of the other soil orders. The wetlands of the tropics present a separate and unique habitat condition and those along the coast have some special soils.

REFERENCES 1. Tootil, E., Ed.; Dictionary of Biology; Intercontinental Kook Productions, Ltd.: Maidenhead, Berkshire, England, 1980. 2. Eswaran, H.; Beinroth, F.H.; Kimble, J.; Cook, T. Soil diversity in the tropics: implications for agricultural development. In Myths and Science of Soils of the Tropics; Lal, R., Sanchez, P.A., Eds.; Soil Sci. Soc. Am. Spec. Publ. 1992; 29, 1–16. 3. Soil survey staff. In Soil Taxonomy: A Basic System of Soil Classification for Making and Interpreting Soil Surveys, 2nd Ed.; Natural Resources Conservation Service. US Department of Agriculture, Handbook 436; US Government Printing Office: Washington, DC, 1999. 4. Keys to Soil Taxonomy, 9th Ed.; US Government Printing Office, 2003.

Global Warming: Carbon Sequestration to Mitigate Keith E. Idso Center for the Study of Carbon Dioxide and Global Change, Tempe, Arizona, U.S.A.

Sherwood E. Idso United States Water Conservation Laboratory, Phoenix, Arizona, U.S.A.

INTRODUCTION Concomitant with mankind’s growing numbers and the progression of the Industrial Revolution, there has been a significant increase in the burning of fossil fuels (coal, gas, and oil) over the past 200 yr, the carbon dioxide emissions from which have led to ever-increasing concentrations of atmospheric CO2. This ‘‘large-scale geophysical experiment,’’ to borrow the words of two of the phenomenon’s early investigators,[1] is still ongoing and expected to continue throughout the current century. Furthermore, this enriching of the air with CO2 is looked upon with great concern, because CO2 is an important greenhouse gas, the augmentation of which is believed by many to have the potential to produce significant global warming. Therefore, and because of perceived serious consequences, such as the melting of polar ice, rising sea levels, coastal flooding, and more frequent and intense droughts, floods, and storms,[2] a concerted effort is underway to slow the rate at which CO2 accumulates in the atmosphere, with the goal of stabilizing its concentration at a level that would prevent dangerous anthropogenic interference with the planet’s climate system. One of the more promising ways of reducing the rate of rise of the air’s CO2 content is to encourage land management policies that promote plant growth, which removes CO2 from the atmosphere and sequesters its carbon, first in vegetative tissues and ultimately in soils. Some of these policies deal with managed forests and agro-ecosystems, while others apply to natural ecosystems, such as unmanaged forests and grasslands. In all instances, however, questions abound. Can carbon inputs to soils really be enhanced or carbon losses reduced? Can carbon storage in recalcitrant fractions of soil organic matter be increased, making it possible to successfully maintain new stores of sequestered carbon for long periods? And what if global warming runs wild? Will the ensuing rise in temperature stimulate plant and microbial respiration rates, returning even more CO2 to the air than is removed by photosynthesis and leading to a negative net ecosystem exchange of carbon? These important questions rank Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001665 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

high on the priority lists of many research organizations concerned about the planet’s future climate and the sustainability of the biosphere.

REMOVING CARBON FROM THE AIR The Role of Man There are only two ways to significantly increase the natural flux of carbon from the atmosphere to the biosphere within the time frame required for effective ameliorative action if the ongoing rise in the air’s CO2 content is indeed a bona fide global warming threat: 1) increase the rate of vegetative CO2 assimilation (photosynthesis) per unit leaf area and=or 2) increase the total plant population of the globe, i.e., leaf area per unit land area. Additionally, these things must be done without increasing the rate at which carbon is lost from the soil. Man can do certain things to promote both of these phenomena while meeting the latter requirement as well. He can, for example, increase the rate of CO2 assimilation per unit leaf area in agro-ecosystems by supplying additional nutrients and water to his crops. As has recently been noted, however, there are significant carbon costs associated with the production and application of fertilizers, as well as the transport of irrigation water; and factoring the CO2 emissions of these activities into the equation often results in little net CO2 removal from the atmosphere via these intensified agricultural interventions.[3] Man can also draw more CO2 out of the air by increasing the acreage of land devoted to growing crops, but this approach simultaneously releases great stores of soil carbon built up over prior centuries. When the plow exposes buried organic matter and it is oxidized, for example, prodigious amounts of CO2 are produced and released to the atmosphere. But if a transition to less intensive tillage is made on fields that have a long history of conventional management and have thus been largely depleted of carbon, there is a good opportunity for nature to rebuild previously lost stores of soil organic matter.[4] 769

770

This approach to carbon sequestration is doubly beneficial for it results in a net removal of CO2 from the atmosphere at the same time that it enhances a whole host of beneficial soil properties.[4,5] Also, abandoned farmlands will gradually replenish their carbon stores, both above- and below-ground, as native vegetation gradually reestablishes itself upon them. And, of course, the process can be hastened and made even more effective if trees are planted on such lands. Even without trees, it has been estimated that agricultural ‘‘best management practices’’ that employ conservation tillage techniques have the potential to boost the current U.S. farm and rangeland soil carbon sequestration rate of 20 million metric tons of carbon per year to fully 200 million metric tons per year,[6] which is approximately 13% of the country’s yearly carbon emissions.[7] Commercial forests also offer excellent opportunities for CO2 removal from the air for considerable periods of time, especially when harvested wood is used to produce products that have long lifetimes. In addition, since some species of trees, such as many of those found in tropical rainforests,[8] can live in excess of a thousand years, CO2 can be removed from the atmosphere and sequestered within their tissues—if man protects the trees from logging—until either long after the Age of Fossil Fuels has run its course or until significant changes in energy systems have reduced our dependence on fossil fuels and the CO2 content of the air has returned to a level no longer considered problematic. Furthermore, carbon transferred to the soil beneath the trees via root exudation and turnover has the potential to remain sequestered even longer.

The Role of Nature The fact that the biosphere has maintained itself over the eons in the face of a vast array of environmental perturbations (albeit with significant modifications) suggests that earth’s plant life has great resiliency and may even be able to exert a restraining influence on climate change. A particularly important negative feedback of this type is the biosphere’s ability to intensify its rate of carbon sequestration in the face of rising atmospheric CO2 concentrations, as this phenomenon slows the rate of rise of the air’s CO2 content and thereby reduces the degree of intensification of the atmosphere’s greenhouse effect. This particular climate-moderating influence of atmospheric CO2 enrichment was first described in quantitative terms by Idso.[9,10] It begins when the aerial fertilization effect produced by the rising CO2 content of the atmosphere elicits an increase in plant CO2 assimilation rate per unit leaf area and when the concomitant plant water use efficiency-enhancing


Global Warming: Carbon Sequestration to Mitigate

effect of the elevated CO2 leads to an increase in the total plant population of the globe, due to the ability of more water-use-efficient plants to live and successfully reproduce in areas where it was formerly too dry for them to survive. In fact, these two effects are so powerful, they may actually be able to stabilize the CO2 content of the atmosphere sometime during the current century, but only if anthropogenic CO2 emission rates do not rise by an inordinate amount in the interim.[9,10] At the very least, together with the things man can do, they have the potential to ‘‘buy time’’ until other less-CO2-emitting technologies become available.[11]

KEEPING CARBON IN THE SOIL As more carbon is added to soils via CO2-enhanced root growth, turnover and exudation, as well as from CO2-induced increases in leaf litter and other decaying plant parts, the trick of significantly augmenting soil carbon sequestration is to keep at least the same percentage of this carbon in the soil as has historically been the case and to do so in the face of potential global warming. A number of studies have addressed various aspects of this subject in recent years, with most of them finding that atmospheric CO2 enrichment has little to no significant effect on plant litter decomposition rates. Furthermore, in nearly all of the cases where elevated CO2 was observed to impact this phenomenon, the extra CO2 was found to actually slow the rate of plant decomposition.[12] Much the same results have been obtained when analogous studies have used temperature as the independent variable. Warming has had either no effect on CO2 evolution from the soil, or it has led to an actual decrease in CO2 loss to the atmosphere.[13] Hence, the balance of evidence obtained from these studies suggests that the same—or a greater—percentage of plant material produced in a world of elevated atmospheric CO2 concentration (and possibly higher mean air temperature) would indeed be retained in the soils of the terrestrial biosphere. Even more compelling are the results of experiments where scientists have made direct measurements of changes in soil carbon storage under conditions of elevated atmospheric CO2. Nearly every such study has observed increases in soil organic matter. In a FreeAir CO2 Enrichment (FACE) experiment where portions of a cotton field were exposed to a 50% increase in atmospheric CO2, for example, Leavitt et al.[14] found that 10% of the organic carbon present in the soil below the CO2-enriched plants at the conclusion of the three-year experiment came from the extra CO2 supplied to the FACE plants. In addition, some


771

Table 1 Potential rates of carbon sequestration (kilograms carbon per hectare per year) due to land management practices that could be employed for this purpose Improved rangeland management

50 to 150

Improved pastureland management Commercial fertilizer applications Manure applications Use of improved plant species

100 to 200 200 to 500 100 to 300

Improved grazing management

300 to 1300

Nitrogen fertilization of mountain meadows

100 to 200

Restoration of eroded soils

50 to 200

Restoration of mined lands

1000 to 3000

Conversion of cropland to pasture

400 to 1200

Conversion of cropland to natural vegetation

600 to 900

Conversion from conventional to conservation tillage No till Mulch till Ridge till

500 500 500

(Adapted from data reported by Follett, R.F.; Kimble, J.M.; Lal, R. The Potential of U.S. Grazing Lands to Sequester Carbon and Mitigate the Greenhouse Effect; Lewis Publishers, Boca Raton, FL, 2001; 1– 442, and by Ref.[4].)

of the stored carbon had made its way into a very recalcitrant portion of the soil organic matter that had an average soil residence time of 2200 yr. Here, too, most experiments indicate that concomitant increases in temperature do not negate the increased carbon storage produced by atmospheric CO2 enrichment. In a two-year study of perennial ryegrass grown at ambient and twice-ambient atmospheric CO2 concentrations, as well as ambient and ambient þ3 C temperature levels, for example, Casella and Soussana[15] determined that the elevated CO2 increased soil carbon storage by 32% and 96% at low and high levels of soil nitrogen supply, respectively, ‘‘with no significant increased temperature effect.’’ Hence, as in the case of studies of plant decomposition rates, the balance of evidence obtained from these studies also suggests that the same—or a greater— percentage of plant material produced in a world of elevated atmospheric CO2 concentration (and possibly higher mean air temperature) would indeed be retained in the soils of the terrestrial biosphere.

CONCLUSIONS As the air’s CO2 content continues to rise, there will almost certainly be a significant upward trend in the


yearly production of terrestrial vegetative biomass, due to the growth-enhancing aerial fertilization effect of atmospheric CO2 enrichment and the concomitant CO2-induced increase in plant water use efficiency that enables plants to grow where it is currently too dry for them. Experimental evidence further suggests that at least the same percentage—but in all likelihood more—of this yearly-increasing mass of plant tissue will be sequestered in earth’s soils. Consequently, it is almost impossible to conclude that the carbon sequestering prowess of the planet will not be greatly enhanced in the years ahead, even without any overt actions on the part of man. Hence, if the nations of the earth were to implement even a modicum of carbon-conserving measures—such as 1)using minimum tillage techniques wherever possible in agricultural settings; 2) allowing abandoned agricultural land to revert to its natural vegetative state; 3)allowing stands of trees that can grow to very old age to actually do so; and 4) employing wise forestry practices to produce wood for making products that have long lifetimes— it is possible that the antiwarming feedback produced by the subsequent removal of CO2 from the atmosphere would be sufficient to keep the risk of potential greenhouse gas-induced global warming at an acceptable level. Estimates of the carbon-sequestering power of some of these ‘‘best management practices’’ are given in Table 1.

REFERENCES 1. Revelle, R.; Suess, H.E. Carbon dioxide exchange between atmosphere and ocean and the question of an increase of atmospheric CO2 during the past decades. Tellus 1957, 9, 18–27. 2. Intergovernmental panel on climate change. In Climate Change 2001: The Scientific Basis, Summary for Policy Makers and Technical Summary of the Working Group I Report; Cambridge University Press: Cambridge, UK, 2001; 1–98. 3. Schlesinger, W.H. Carbon sequestration in soils: some cautions amidst optimism. Agric. Ecosys. Environ. 2000, 82, 121–127. 4. Lal, R.; Kimble, J.M.; Follett, R.F.; Cole, C.V. The Potential for U.S. Cropland to Sequester Carbon and Mitigate the Greenhouse Effect; Sleeping Bear Press: Chelsea, MI, 1998; 1–128. 5. Idso, S.B. Carbon Dioxide and Global Change: Earth in Transition; IBR Press: Tempe, AZ, 1989; 1–292. 6. Jawson, M.D.; Shafer, S.R. Carbon credits on the Chicago board of trade? Agric. Res. 2001, 49 (2), 2. 7. Comis, D.; Becker, H.; Stelljes, K.B. Depositing carbon in the bank. Agric. Res. 2001, 49 (2), 4–7. 8. Chambers, J.Q.; Higuchi, N.; Schimel, J.P. Ancient trees in Amazonia. Nature 1998, 391, 135–136. 9. Idso, S.B. The aerial fertilization effect of CO2 and its implications for global carbon cycling and maximum

772

greenhouse warming. Bull. Amer. Meteorol. Soc. 1991, 72, 962–965. 10. Idso, S.B. Reply to comments of L.D. Danny Harvey, Bert Bolin, and P. Lehmann. Bull. Amer. Meteorol. Soc. 1991, 72, 1910–1914.. 11. Izaurralde, R.C.; Rosenberg, N.J.; Lal, R. Mitigation of climatic change by soil carbon sequestration: issues of science, monitoring, and degraded lands. Adv. Agron. 2001, 70, 1–75. 12. Nitschelm, J.J.; Luscher, A.; Hartwig, U.A.; van Kessel, C. Using stable isotopes to determine soil carbon input differences under ambient and elevated atmospheric CO2 conditions. Global Change Biol. 1997, 3, 411–416.



13. van Ginkel, J.H.; Whitmore, A.P.; Gorissen, A. Lolium perenne grasslands may function as a sink for atmospheric carbon dioxide. J. Environ. Quality 1999, 28, 1580–1584. 14. Leavitt, S.W.; Paul, E.A.; Kimball, B.A.; Hendrey, G.R.; Mauney, J.R.; Rauschkolb, R.; Rogers, H.; Lewin, K.F.; Nagy, J.; Pinter, P.J., Jr.; Johnson, H.B., Jr. Carbon isotope dynamics of free-air CO2-enriched cotton and soils. Agric. For. Meteorol. 1994, 70, 87–101. 15. Casella, E.; Soussana, J.-F. Long-term effects of CO2 enrichment and temperature increase on the carbon balance of a temperate grass sward. J. Exp. Bot. 1997, 48, 1309–1321.

Grass Strip Hydrology Hossein Ghadiri Calvin W. Rose Faculty of Environmental Sciences, Griffith University, Nathan Campus, Nathan, Queensland, Australia

INTRODUCTION The use of vegetated strips is one of the simplest and most widely employed methods of reducing the flux of eroding soil and associated chemicals across slopes (Fig. 1) and into streams. Grass strips of various type, width, spacing, density, rigidity, and strength have been used for soil erosion control on sloping lands and for water quality control in riparian zones with varying degrees of success.[1–3] It was widely believed, with some support from field studies, that the effectiveness of buffer strips in sediment trapping was largely because of a filter-type action, with strips filtering out the suspended sediment as runoff passes through them, leaving the emerging runoff significantly cleaner in terms of sediment, nutrients, and soil-sorbed contaminants. Because of such perceived mode of operation, buffer strips have been commonly referred to as ‘‘filter strips.’’[4,5] However, most recent studies have shown that a filtering action is not a major contributing process to buffer strip effectiveness. Considerable experimental evidence in the literature suggests little or no net deposition within the buffer strips.[6–9] Some have reported the occurrence of erosion, rather than deposition, inside the grass strips.[4,8] Recent studies have sought to understand the mechanics of flow through grass strips, seeking the real reason for the inconsistencies observed in their effectiveness.[2,7–11] It has been recognized that a flowresistive element such as a cross-slope strip of vegetation modifies the hydrology of overland flow, and that this modification has implications for the transport and deposition of sediment and associated nutrients in and around the strips.[7,8,11,12] It is the hydraulic consequence for overland flow when it meets and flows through a resistive element, which first needs to be understood to ascertain the effectiveness, or otherwise, of grass strips in reducing erosion, enhancing deposition, and reducing the transportation of pollutants into surface water bodies. Such understanding is also necessary for the development of models and the prediction of buffer strip effectiveness under different conditions. This article covers the most recent progresses made in this area. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120014316 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

EFFECT OF BUFFER STRIPS ON FLOW HYDRAULICS One of the main hydrologic features of flow through grass strips is the formation of a pronounced backwater,[8] or a zone of hydraulic adjustment.[9,13] The length of this zone appears to be dependent on land slope, strip density, and flow rate (Fig. 2). At any constant flow rate, the length of backwater B has been shown to be linearly related to the ratio of strip density over slope:

B ¼ k

D S

ð1Þ

where k is a proportionality constant, D is percent grass coverage in the strip, and S is slope. Eq. (1) shows that decreasing density and increasing slope both have the effect of decreasing backwater length[8] (Fig. 3). The starting point of the backwater region is a hydraulic jump, which moves closer to the strip as the slope increases or as the grass density decreases. The front edge of the backwater is usually sharp and clear at high slopes (Figs. 2 and 3), getting less distinct as it moves away from the strips with decreasing slope or increasing grass density. At low slopes of around 1%, the front edge of the backwater breaks into a number of waves. The width of the strips does not appear to play an important role in backwater formation or its characteristics.[8] Theoretical interpretations of flow through grass strips are very few. However, there is a long history of investigation of hydraulic jumps in the hydraulic literature. Chow[13] developed the hydraulic jump theory for deep rectangular channel flow encountering solid barriers such as dams and wears. This interpretation does not apply to our condition of shallow unconfined flow being partially blocked by porous barriers such as grass or nail strips. Rose et al.[9] have shown that the change in hydraulics because of the presence of a buffer strip can be understood in terms of the conservation of momentum theory. 773

774

Grass Strip Hydrology

Fig. 1 Strips of Vetiver grass used as a soil erosion control technique in Queensland, Australia. (Courtesy of Mr. Cyril Ciesiolka of DPI, Toowoomba, Queensland, Australia.) (View this art in color at www.dekker. com.)

EFFECT OF BUFFER STRIPS ON SEDIMENT TRANSPORT Runoff sediment concentration results from the balance of two quite rapid processes: entrainment (or re-entrainment) by the overland flow, and deposition. Even a modest reduction in flow velocity in the backwater region produced by the vegetation buffer strips reduces the entrainment rate, resulting in net deposition in this region. Thus the effect of grass strips on erosion=sedimentation is mainly through their impact on surface hydrology.

Net deposition of coarse sediment load commences at the starting edge of the backwater (or at the hydraulic jump) (Fig. 2). This site of initial net deposition approaches the strip with an increase in slope, or a decrease in strip density, eventually burying and entering the strip. Sediment deposition generally follows the hydrologic pattern established at the early stages of the flow, as predicted by Eq. (1). Little, if any, deposition appears to take place inside the strips prior to the point of grass collapse.[6,8,14] Sediment that deposits in the backwater moves into the strip zone only after burying the first few rows of the vegetation.[2] The unburied rows then become the new front edge of the strip. However, once the first rows of the strips give way, the process usually continues until the entire width of the strip collapses at one point or more (Fig. 4). Many researchers have reported that the width of the grass strip has little effect on the efficiency of the strips in slowing down the flow, or unloading its sediment content.[2,15] However, there are others who have reported the opposite.[1,3] The disagreement between the two groups is yet to be resolved. THEORETICAL INTERPRETATION OF SEDIMENT TRANSPORT THROUGH BUFFER STRIPS

Fig. 2 Illustrating flow and sediment deposition in the backwater region of a barrier strip.


Mass conservation of deposited sediment was used by several researchers[16,17] to describe the resulting change in the shape of the deposited layer. However, their models do not represent the hydrology and net deposition upslope of the grass strips (Fig. 2), and only deal with single-size particles. A more appropriate theoretical


775

Fig. 3 Effect of slope on the length of backwater (flow height recorded digitally; flow is from left to right and the two long vertical lines on each trace show the beginning and the end of the grass strip).

explanation of flow through porous barriers is given by Rose et al.,[12] who showed that the appropriate equation to solve (under steady flow conditions) is as follows:

class i, x is downslope distance, di the rate of deposition per unit area, and rri is the rate of re-entrainment of newly deposited sediment of size class i.

d dci ¼ di rri Mdi ¼ q dx dt

EFFECT OF BUFFER STRIPS ON CHEMICAL TRANSPORT

where t is time, Mdi is the mass per unit area of sediment of size class i in the region of net deposition, q is volumetric water flux, ci is sediment concentration of size

Soil chemicals are preferentially sorbed to the finer fractions of soil.[18,19] Buffer strip influences the transport of sorbed chemicals mainly through sediment sorting

Fig. 4 Collapsed grass strip on 8% slope (flow from left to right; total collapse happened after the first few rows of grass were buried under the deposited sediment; vertical white bar is a 15-cm ruler). (View this art in color at www.dekker.com.)


776

processes, which take place in the backwater. Sediment passing though the strips is significantly finer than that initially dislodged by the flow. It was shown by Ghadiri and Rose[19] that the concentrations of organic matter, sorbed nutrients, and agricultural chemicals can be significantly higher on the finer particles. The preferential deposition of larger aggregates and particles in the backwater leaves the sediment, which emerges from the buffer strip, relatively enriched in fine particles and thus in soil-sorbed chemicals. Such transmitted fine particles, together with their chemical load, may stay in suspension until entry into receiving waters. Therefore, grass strips are less effective in reducing overland transport of solutes or solids-associated chemical pollutants than they are in reducing sediment load.

CONCLUSIONS Buffer strips behave like porous barriers against flow, creating backwater regions upslope of the strips with reduced flow velocity and increased depth in a length that varies with slope, strip density, and flow rate. For a constant flow rate, there is a linear relationship between backwater length and the ratio of strip density over slope. Momentum theory appears to provide a reasonably good prediction, both of the shape of the water profile within the strips, and of the slope and extent of backwater region. The zone of hydraulic adjustment or backwater is a zone of considerable deposition from sediment-laden flow. The efficiency of the grass strips in slowing down the flow and unloading its sediment in the backwater region appears to be largely independent of the width of strips in the flow direction. The layer of deposited sediment formed in the backwater region tends to be richer in larger particles than the eroding sediment. Sediment size distribution appears to be a dominant factor governing the efficiency of the buffer strip in trapping sediment. Soil-sorbed nutrients and other agricultural chemicals are mainly attached to finer soil particles, which are more likely to pass through the strips largely unchanged. Thus the sediment that emerges from a buffer strip can be enriched in such chemicals relative to the eroding soil. Hence although buffer strips can be very efficient in forcing the deposition of suspended sediment in the backwater region on certain slopes, they are helpful, but less effective, in preventing chemical pollutants from entering surface water bodies.

REFERENCES 1. Dickey, E.C.; Vanderholm, D.H. Vegetation filter treatment of livestock feedlot runoff. J. Environ. Qual. 1981, 10, 279–284.



2. Dillaha, T.A.; Reneau, R.B.; Mostaghimi, S.; Lee, D. Vegetation filter strips for agricultural nonpoint source pollution control. Trans. ASAE 1989, 32, 513–519. 3. Van Dijk, P.M.; Kwaad, F.J.P.; Klapwijk, M. Retention of water and sediment by grass strips. Hydrol. Process. 1996, 10, 1069–1080. 4. Loch, L.J.; Espigares, T.; Costantini, A.; Garthe, R.; Bubb, K. Vegetative filter strips to control sediment movement in forest plantations: validation of a simple model using field data. Aust. J. Soil Res. 1999, 37, 929–946. 5. Jin, C.X.; Romkens, M.J.M. Experimental studies of factors in determining sediment trapping in vegetative filter strips. Trans. ASAE 2001, 44, 277–288. 6. Lidgi, E.; Morgan, R.P.C. Contour grass strips: a laboratory simulation of their role in soil erosion control. Soil Technol. 1995, 8, 109–117. 7. Ghadiri, H.; Hogarth, W.; Rose, C.W. The Effectiveness of Grass Strips for the Control of Sediment and Associated Pollutant Transport of Runoff; Stone, M., Ed.; International Association of Hydrologic Sciences (IAHS) Publication No. 263, 2000; 83–91. 8. Ghadiri, H.; Rose, C.W.; Hogarth, W.L. The influence of grass and porous barrier strips on runoff hydrology and sediment transport. Trans. ASAE 2001, 44, 259–268. 9. Rose, C.W.; Hogarth, W.; Ghadiri, H.; Parlange, J.-Y.; Okom, A. Overland flow to and through a segment of uniform resistance. J. Hydrol. 2002, 255, 134–150. 10. Hairsine, P.B. In Comparing Grass Filter Strips and Near-Natural Riparian Forests for Buffering Intense Hillslope, Sediment Sources, Proceedings of the 1st National Conference on Stream Management in Australia; CRC for Catchment Hydrology, Monash University: Melbourne, Australia, 1996. 11. Dabney, S.M.; Meyer, L.D.; Harmon, W.C.; Alonso, C.V.; Foster, G.R. Depositional patterns of sediment trapped by grass hedges. Trans. ASAE 1995, 38, 1719–1729. 12. Rose, C.W.; Yu, B.; Hogarth, W.; Okom, E.A.; Ghadiri, H. Theoretical interpretation of the spatial and size distribution of sediment deposited by buffer strips from flow at modest land slopes. J. Hydrol. 2003, in press. 13. Chow, V.T. Open Channel Hydraulics; McGraw-Hill Book Company: Singapore, 1959. 14. Boubakari, M.; Morgan, R.P.C. Contour grass strips for soil erosion control on steep lands: a laboratory evaluation. Soil Use Manage. 1999, 15, 21–26. 15. Smith, C.M. Riparian afforestation effects on water yields and water quality in pasture catchments. J. Environ. Qual. 1992, 21, 237–245. 16. Munoz-Carpena, R.; Pearson, J.E.; Gilliam, J.W. Modelling hydrology and sediment transport in vegetative filter strips. J. Hydrol. 1999, 214, 111–129. 17. Deletic, A. Modelling of water and sediment transport over grasses areas. J. Hydrol. 2001, 248, 168–182. 18. Ghadiri, H.; Rose, C.W. Sorbed chemical transport in overland flow: Part 1. An enrichment mechanism for sorbed nutrients and pesticides. J. Environ. Qual. 1991, 20, 628–633. 19. Ghadiri, H.; Rose, C.W. Sorbed chemical transport in overland flow: Part 2. Enrichment ratio of sorbed chemicals and its variation with time, particle size and erosion process. J. Environ. Qual. 1991, 20, 634–641.

Grasslands Soils Douglas D. Malo South Dakota State University, Brookings, South Dakota, U.S.A.

INTRODUCTION Soil science developed in recent times in response to problems. In Europe during the 1800s, food shortages, social upheavals, and declining soil productivity brought about the need to study soil to improve and increase its productivity. At the same time, in Russia, a need arose to administer and manage geographically diverse soil resources. Russia had large areas of fertile, productive soils unlike those in Europe. As a result, Russian scientists developed an inventory of agricultural resources and determined the factors causing soils to vary across Russia. Soils were found to have relationships with climatic and vegetative zones.[1] It is from these concepts that the living soil individual was developed. Soil is a natural body and not a geologic formation. With time, soil develops from the parent material under the influence of the climate, vegetation, and topography (relief ). These five factors of soil formation are interdependent and not independent. Changing one soil-forming factor often changes other soil-forming factors. Changing any one, some, or all of the soil-forming factors causes differences in soils and soil profiles.[2] Soils and plants interact and evolve together forming different ecosystems. These soil–plant relationships are often most strongly expressed when native vegetation is present. The soils in grasslands, a major ecosystem, are very different from other soils due to this interaction.

the steppe regions of Europe, Russia, Mongolia, and Northern China) are dominated by grassland soils (Fig. 1). The largest grassland area in the world is found in Kazakhstan, Russia, and the Ukraine. Generally, small grains and grain sorghum are raised in the drier grassland regions. The warmer, humid grasslands are better suited for row crops like maize (corn) and soybeans. Where slopes are too great for cultivation or the climate is not favorable, grassland soils are used for pasture and rangeland. Native grassland types include: desert grasslands, 75,000 yr, A-Bkm horizons may form and persist in arid to semiarid climates with well-developed A-E-Bt horizons forming in subhumid and humid climates.

CONCLUSIONS Knowing the formation and distribution of SIC components is important for the mapping and classification of soils, understanding carbonate equilibria reactions, better interpreting of the role of soils in the global carbon cycle, and understanding feedback mechanisms that soil may provide in a future greenhouse world.

REFERENCES 1. Birkeland, P.W. Soils and Geomorphology; Oxford University Press: New York, 1999. 2. Monger, H.C.; Daugherty, L.A.; Lindemann, W.C.; Liddell, C.M. Microbial precipitation of pedogenic calcite. Geology 1991, 19, 997–1000. 3. Nordt, L.C.; Hallmark, C.T.; Wilding, L.P.; Boutton, T.W. Quantifying pedogenic carbonate accumulations using stable carbon isotopes. Geoderma 1998, 82, 115–136. 4. Schlesinger, W.H. Biogeochemistry: Analysis of Global Change, 1st Ed.; Academic Press: San Diego, California, 1991; 1–443. 5. Amundson, R.G.; Chadwick, O.A.; Sowers, J.M. A comparison of soil climate and biological activity along an elevation gradient in the eastern mojave desert. Oecologia 1989, 80, 395–400. 6. Cerling, T.E. The stable isotopic composition of modern soil carbonate and its relationship to climate. Earth and Planetary Science Letters 1984, 71, 229–240. 7. Suarez, D.L. Impact of agriculture on CO2 as affected by changes in inorganic carbon. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, Florida, 2000; 257–272. 8. Schwab, A.P. The soil solution. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, Florida, 2000; B85–B120. 9. Nordt, L.C.; Wilding, L.P.; Drees, L.R. Pedogenic carbonate transformations in leaching soil systems: implications for the global C cycle. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, Florida, 2000; 43–64.

Inorganic Carbon: Climate and Time

10. Retallack, G.J. The environmental factor approach to the interpretation of paleosols. In Factors of Soil Formation: A Fiftieth Anniversary Retrospective; Amundson, R., Harden, J., Singer, M., Eds.; Special Publication Number 33; Soil Science Society of America: Madison, Wisconsin, 1994; 31–64. 11. Royer, D.L. Depth to pedogenic carbonate as a paleoeprecipitation Indicator? Geology 1999, 27, 1123–1126. 12. Gile, L.H.; Hawley, J.W.; Grossman, R.B. Soils and Geomorphology in the Basin and Range Area of Southern New Mexico-Guidebook to the Desert Project;


885

Memoir 39; New Mexico Bureau of Mines and Mineral Resources: Socorro, New Mexico, 1981; 1–222. 13. Machette, M.N. Calcic soils of the Southwestern United States. In Soils and Quaternary Geomorphology of the Southwestern United States; Weide, D.L., Ed.; Geological Society of America: Boulder, Colorado, 1985; 1–21. 14. Rabenhorst, M.C.; West, L.T.; Wilding, L.P. Genesis of calcic and petrocalcic horizons in soils over carbonate rocks. In Occurrence, Characteristics, and Genesis of Carbonate, Gypsum, and Silica Accumulations in Soils; Nettleton, W.D., Ed.; Soil Science Society of America: Madison, Wisconsin, 1991; 61–74.

Inorganic Carbon: Composition and Formation H. Curtis Monger New Mexico State University, Las Cruces, New Mexico, U.S.A.

Larry P. Wilding Texas A&M University, College Station, Texas, U.S.A.

INTRODUCTION

Cations, such as Ca2þ, Mg2þ, Fe2þ, Mn2þ, and Naþ, precipitate with HCO3 (which is the dominant anion between pH 6.5 and 10.5) and CO32 (which is the dominant anion above pH 10.5) to form a variety of carbonate minerals. The reaction of Ca(aq)2þ with HCO3(aq) to form calcite is illustrated below:

Individual carbonate crystals of pedogenic origin are generally too small to be seen with the unaided eye. Yet when concentrated together, their combined presence takes on a white color with a variety of macroscopic forms. These forms include carbonate filaments (also called mycelia, pseudomycelia, and threads), films, coatings, soft spheroidal segregations (white eyes), nodules, cylindroids, concretions, glaebules, and veins. Soil fabric which is impregnated with carbonate to the point that it occurs as an essentially continuous medium has been termed ‘‘k-fabric.’’[4] Stages of morphogenetic carbonate accumulation, in which progressively greater amounts of carbonate occur in progressively older soils, are important chronologic indicators. The calcic and petrocalcic horizon are diagnostic horizons in Soil Taxonomy.[5] Calcic horizons generally contain greater than 15% carbonate by weight. Petrocalcic horizons are indurated forms of calcic horizons. Examples of these horizons are shown in Figs. 1A and B. Examples of carbonate crystals as viewed with optical microscopy and scanning electron microscopy are shown in Figs. 1C and D. Dissolution of carbonates in soil systems can be represented by the following reaction (Eq. 3). In humid regions, soluble products of this weathering reaction flux through the vadose zone into groundwater, or precipitate as pedogenic carbonates deep in the soil or geologic system. In arid regions, soluble products precipitate at relatively shallow depths as a result of sparse rainfall and insufficient leaching.

CaðaqÞ 2þ þ 2HCO3ðaqÞ

HCO3ðaqÞ þ HðaqÞ þ þ CaCO3ðsÞ

It has become increasingly common for ‘‘soil inorganic carbon’’ to mean soil carbonate mineral carbon, mainly CaCO3. In the strict sense, however, inorganic carbon not only encompasses carbon in carbonate minerals, but also carbon in the carbonic acid system.[1] The carbonic acid system includes gaseous carbon dioxide (CO2(g)), aqueous carbon dioxide (CO2(aq)), carbonic acid (H2CO3(aq)), bicarbonate ion (HCO3(aq)), and carbonate ion (CO3(aq)2).

COMPOSITION In the soil solution, as with other solutions, the interaction of these species can be represented by the following reaction:[1] CO2ðgÞ þ H2 OðlÞ ¼ H2 CO3ðaqÞ ¼ HCO3ðaqÞ þ HðaqÞ þ ¼ CO3ðaqÞ 2 þ 2HðaqÞ þ

¼ CaCO3ðsÞ þ CO2ðgÞ þ H2 OðlÞ

ð1Þ

ð2Þ

There are about 60 carbonate minerals, which in addition to calcite, include aragonite (CaCO3), dolomite [CaMg(CO3)2], siderite (FeCO3), magnesite (MgCO3), rhodocrosite (MnCO3), cerussite (PbCO3), and malachite [CuCO3Cu(OH)2]. In soil, the overwhelmingly abundant carbonate mineral is calcite.[2] In unique soil environments, however, other carbonate minerals have been found, such as pedogenic siderite and dolomite.[3] 886 Copyright © 2006 by Taylor & Francis

¼ CaðHCO3 Þ2ðaqÞ

ð3Þ

FORMATION Being located in an arid, semiarid, or subhumid climate is the primary factor that controls carbonate formation. In many areas, the boundary between carbonateaccumulating soil and noncarbonate-accumulating soil is about 500 mm (20 in.) mean annual rainfall.[6] This Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001712 Copyright # 2006 by Taylor & Francis. All rights reserved.

Inorganic Carbon: Composition and Formation

887

Fig. 1 Examples of inorganic carbon as it exists in the field and under the microscope. The white horizon is a calcic horizon in (A), and petrocalcic horizon in (B). The small golden crystals in (C) are calcite crystals coating sand grains; the black region is a pore space as it appears in cross-polarized light. A calcified fungal filament (cf ) viewed with scanning electron microscopy is shown in (D).

relationship is confounded, however, by the effects of soil temperature, soil drainage, nature and properties of the parent material (e.g., soil texture, carbonate content, carbonate mineralogy, and porosity), soil drainage, landform position, geomorphic stability, and effectiveness of precipitation (rainfall intensity and duration). Hence, there are many examples in humid and subhumid environments where soil carbonate persists in the soil system at depths inconsistent with regional models. In humid regions, for example, inorganic carbon persists as calcite or dolomite detritus in soils derived from certain parent materials (e.g., calcareous loess, till, outwash, alluvial deposits, sedimentary and metamorphic rocks). In seasonally wet soils, carbonate can accumulate in upper subsoils from capillary rise of bicarbonates via evaporative pumping from shallow groundwater.[7] In addition, carbonate minerals can occur in wetland soils which commonly contain soluble carbonates, bicarbonates or carbonic acid depending on the pH of the local environment.


Pedogenic vs. Geogenic Carbonate Many soils develop in calcareous parent materials. For these soils, it has been a challenge quantifying carbonate that formed in the soil profile vs. carbonate mechanically inherited from parent material. Carbonate formed in the soil profile has been termed ‘‘secondary,’’ ‘‘authigenic,’’ or ‘‘pedogenic.’’[4,8] On the other hand, carbonate mechanically inherited from parent material has been termed ‘‘primary,’’ ‘‘geogenic,’’ or ‘‘lithogenic.’’[2,9] Criteria for distinguishing pedogenic from geogenic carbonates involve the scrutiny of both field and laboratory evidence. Field evidence, for example, includes differences in the presence of marine fossils, carbonate morphology (such as nodules, pendants, and laminar caps which indicate pedogenic), and distribution patterns with depth, where, for example, a carbonate horizon of pedogenic origin is overlain and underlain by noncalcareous soil. Laboratory evidence includes comparing mineralogy, particle size, microfabric, and 13C=12C ratios of

888

carbonate with unknown origin to those of carbonate with known geogenic origin.[10–12] Models of Carbonate Formation There are several processes that cause carbonates to form in soil. Excluding geologic processes, such as lacustrine and deep groundwater cementation that preserves the original sedimentary structure, the formation pedogenic carbonate can broadly be placed into four models—per descensum, per ascensum, in situ, and biogenic models. The per descensum model The per descensum model accounts for carbonate formation resulting from downward moving meteoric water and can be subdivided into three types. First is the dissolution of pre-existing carbonates in the upper profile, their vertical translocation, and their precipitation in the subsoil. This model was invoked to explain why progressively shallower carbonates occur in progressively drier climates.[13] Later, this per descensum model was used as the basis for calculating the number of wetting-fronts required to leach carbonates to a particular depth.[14] In both cases, it was assumed that carbonate was uniformly distributed in parent material at the beginning of pedogenesis. Second is the case in which pedogenic carbonate forms in soils with noncalcareous parent materials. Unlike the model described before, noncalcareous parent material does not have carbonate uniformly distributed throughout the profile at the beginning of pedogenesis. In southern New Mexico, for example, prominent calcic and petrocalcic horizons occur in soils with rhyolite alluvium as parent material. This alluvium would yield low amounts of calcium if the rhyolite particles were thoroughly decomposed, which they were not.[15] Therefore, atmospheric additions, another per descensum model, was judged to be the source of carbonates.[15] Initially calcareous dust was measured and considered to be the source of carbonate. Later it was realized that Ca2þ in rain was an additional, and even larger source of Ca2þ for reaction with soil HCO3 to form carbonates.[15] Building on these per descensum concepts, compartmental models have been constructed that compute the depth, amount, and distribution of pedogenic carbonate as a function of climate and time.[16,17] Third, in addition to vertical illuviation within a soil profile, lateral, downslope migration of the soil solution containing soluble products of carbonate is another per descensum model. In this case, carbonate is thought to precipitate after carbonate-charged waters migrate from upslope positions to lower landscape positions.[18]



The per ascensum model The per ascensum model accounts for carbonate formation resulting from bottom–up movement. A primary example is the capillary rise of Ca2þ and bicarbonate from shallow water tables by evaporative pumping, which leads to the precipitation of carbonates in the upper subsoil.[7] Moreover, chemical studies have shown that in some environments plants promote carbonate formation by transporting Ca2þ upward to the land surface from subsoil, rock, and groundwater.[19] The in situ model Third, in the in situ model, pedogenic carbonate is the result of in-place dissolution and reprecipitation of bedrock composed of marine carbonate.[20] Limestone, for instance, is progressively transformed into pedogenic carbonate as a result of short-range carbonate dissolution and reprecipitation proximal to the depth of the upper contact with limestone. This is a rather unique method to form pedogenic carbonates where the total carbonate content of the zone of enriched pedogenic products is less than the carbonate content of the limestone originally. These pedogenic zones have a much higher macro- and micro-porosity than the limestone. In addition to marine carbonates, the in situ model also includes carbonate formation resulting from inplace chemical weathering of Ca-bearing igneous rock. Upon release into the soil solution by weathering, Ca2þ precipitates with bicarbonate formed from the reaction of water with CO2 generated by root and microbial respiration. In many cases, however, igneous parent material has been considered as an insufficient source of Ca2þ and hence external sources, such as atmospheric additions of Ca2þ, have been sought.[6] The biogenic model Fourth, some plants, microorganisms, and termites produce calcium carbonate. Evidence that various plants play a direct role in carbonate formation comes from the presence of euhedral calcite crystals on plant roots.[21] Moreover, several references in the Russian literature note carbonate formation by plant tissue and the downward translocation of these carbonates with wetting fronts.[22] Evidence that some microorganisms precipitate carbonates is based on observations of calcified bacteria and fungal hyphae with electron microscopy and in vitro laboratory experiments.[23,24] Evidence that termites precipitate carbonate in certain environments is based on the studies of termite mounds in Africa and southeast Asia.[25] Such mounds can be calcareous even though surrounding soils are noncalcareous, making the mounds attractive to native farmers who spread them over their agricultural fields.[25]


CONCLUSIONS The formation of pedogenic carbonate may be dominated by one of the models listed before or may involve several of the models working together in different magnitudes. Understanding pedogenic carbonate formation has been extremely useful for understanding relative ages of geomorphic surfaces and landscape evolution.[15] A knowledge of pedogenic carbonate formation has also been useful for soil classification. Marbut,[26] for instance, used the presence of carbonate as a criterion for the highest category of his soil classification—Pedocal (soils with carbonate accumulation) and Pedalfers (soil with Al and Fe accumulation). Today, studies of pedogenic carbonate have expanded to include questions about paleoclimate, paleoecology, paleoatmospheric composition, global carbon cycles and the greenhouse effect.

REFERENCES 1. Morse, J.W.; Mackenzie, F.T. Geochemistry of Sedimentary Carbonates. Developments in Sedimentology 48; Elsevier: Amsterdam, 1990. 2. Doner, H.E.; Lynn, W.C. Carbonate, halide, sulfate, and sulfide minerals. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, WI, 1989; 279–330. 3. Capo, R.C.; Whipkey, C.E.; Blache`re, J.R.; Chadwick, O.A. Pedogenic origin of dolomite in a basaltic weathering profile, Kohala Peninsula, Hawaii. Geology 2000, 28, 271–274. 4. Gile, L.H.; Peterson, F.F.; Grossman, R.B. The K horizon—a master soil horizon of carbonate accumulation. Soil Sci. 1965, 99, 74–82. 5. Soil Survey Staff. Soil Taxonomy—A Basic System of Soil Classification for Making and Interpreting Soil Surveys, 2nd Ed.; USDA Agriculture Handbook Number 436; U.S. Govt. Printing Office: Washington, DC, 1999. 6. Birkeland, P.W. Soils and Geomorphology, 3rd Ed.; Oxford University Press: New York, 1999; 430. 7. Sobecki, T.M.; Wilding, L.P. Formation of calcic and argillic horizons in selected soils of the texas coast prairie. Soil Sci. Soc. Am. J. 1983, 47, 707–715. 8. Pal, D.K.; Dasog, G.S.; Vadivelu, S.; Ahuja, R.L.; Bhattacharyya, T. Secondary calcium carbonate in soils of arid and semiarid regions of India. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: London, 2000; 149–185. 9. West, L.T.; Wilding, L.R.; Rabenhorst, M.C. Differentiation of pedogenic and lithogenic carbonate forms in texas. Geoderma 1987, 43, 271–287. 10. Rabenhorst, M.C.; Wilding, L.P.; West, L.T. Identification of pedogenic carbonates using stable carbon isotopes and microfabric analysis. Soil Sci. Soc. Am. J. 1984, 48, 125–132.


889

11. Drees, L.R.; Wilding, L.P. Micromorphic record and interpretations of carbonate forms in the rolling plains of texas. Geoderma 1987, 40, 157–175. 12. Nordt, L.C.; Hallmark, C.T.; Wilding, L.P.; Boutton, T.W. Quantifying pedogenic carbonate accumulations using stable carbon isotopes. Geoderma 1998, 82, 115–136. 13. Jenny, H.; Leonard, C.D. Functional relationships between soil properties and rainfall. Soil Sci. 1934, 38, 363–381. 14. Arkley, R.J. Calculation of carbonate and water movement in soil from climatic data. Soil Sci. 1963, 96, 239–248. 15. Gile, L.H.; Hawley, J.W.; Grossman, R.B. Soils and Geomorphology in the Basin and Range Area of Southern New Mexico—Guidebook to the Desert Project; New Mexico Bureau of Mines and Mineral Resources, Memoir 39: Socorro, New Mexico, 1981; 222 pp. 16. McFadden, L.D.; Tinsley, J.C. Rate and depth of pedogenic-carbonate accumulation in soils: formation and testing of a compartment model. In Soils and Quaternary Geology of the Southwestern United States; Weide, D.L., Ed.; Special Paper 203; Geological Society of America: Boulder, CO, 1985; 23–42. 17. Marion, G.M.; Schlesinger, W.H. Quantitative modeling of soil forming processes in deserts: the CALDEP and CALGYP models. In Quantitative Modeling of Soil-Forming Processes; Bryant, R.B., Arnold, R.W., Eds.; Soil Sci. Soc. Am. Spec. Publ. 39: Madison, WI, 1994; 129–145. 18. Scharpenseel, H.W.; Mtimet, A.; Freytag, J. Soil inorganic carbon and global change. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: London, 2000; 27–42. 19. Goudie, A. Duricrusts in Tropical and Subtropical Landscapes; Oxford Univ. Press: London, 1973; 174. 20. Rabenhorst, M.C.; Wilding, L.P. Pedogenesis on the Edwards plateau, Texas: III. A new model for the formation of petrocalcic horizons. Soil Sci. Soc. Am. J. 1986, 50, 693–699. 21. Monger, H.C.; Gallegos, R.A. Biotic and abiotic processes and rates of pedogenic carbonate accumulation in the southwestern United States—relationship to atmospheric CO2 sequestration. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: London, 2000; 273–290. 22. Labova, E. Soils of the Desert Zone of the USSR; Israel Program for Scientific Translation: Jerusalem, Israel, 1967. 23. Phillips, S.E.; Milnes, A.R.; Foster, R.C. Calcified filaments: an example of geological influences in the formation of calcretes in South Australia. Aust. J. Soil Res. 1987, 25, 405–428. 24. Monger, H.C.; Daugherty, L.A.; Lindemann, W.C.; Liddell, C.M. Microbial precipitation of pedogenic calcite. Geology 1991, 19, 997–1000. 25. Thorp, J. Effects of certain animals that live in soils. Sci. Monthly 1949, 68, 180–191. 26. Marbut, C.F. Subcommission II. Classification, nomenclature, and mapping of soils in the United States. Soil Sci. 1928, 25, 61–71.

Inorganic Carbon: Global Stocks Larry P. Wilding Texas A&M University, College Station, Texas, U.S.A.

Lee C. Nordt Baylor University, Waco, Texas, U.S.A.

John M. Kimble United States Department of Agriculture (USDA), Lincoln, Nebraska, U.S.A.

INTRODUCTION In recent years more work has focused on the role of soil organic carbon (SOC),[1–5] than soil inorganic carbon (SIC)[6,7] in the global carbon cycle. One of the constraints to quantifying SIC stocks is the incomplete global dataset. Thus, we can estimate global SIC stocks, but with considerable uncertainty. Further, there is difficulty differentiating lithogenic from pedogenic carbonates because most routine chemical methods determine the total quantity of carbonates in a soil sample without regard to origin. The amount of pedogenic carbonate formed in soils is typically estimated from visible segregations in the field, from chemical analysis in the laboratory under the assumption that the carbonate is pedogenic,[8] or from stable carbon isotope analysis to differentiate lithogenic from disseminated pedogenic carbonate components.[9] Gile and Grossman[10] consider that much of the calcium in carbonates results from carbonate dusts or from calcium in rainwater. Chadwick and Capo[11] believe that at least 95% of the Ca in carbonates of arid and semiarid regions is derived from dusts. Many soils also have lithogenic carbonate inherited from alluvial, till, and bedrock parent materials. It is possible that as little as 10% of the total SIC reported in soils is formed from the authigenic weathering of calcium-bearing minerals. Perhaps even less well understood are the flux rates of SIC stocks (e.g., soluble carbonates, bicarbonates, and carbonic acid) from soil systems into the vadose zone, geologic substrata, ground water aquifers, rivers, lakes, and oceans. The flux rates of inorganic carbon stocks require knowledge of soil chronology coupled with mass balance reconstruction analyses to determine the magnitude of gains or losses.[12,13] This long-term transitory SIC stock is believed to have important impacts on soil carbon sequestration over geologic time periods up to 10 times longer than the SOC stocks.[12–14] Also, the biogenesis of SIC stocks by various organisms (e.g., plant roots, blue green 890 Copyright © 2006 by Taylor & Francis

algae, termites, plankton, etc.), while well documented, is not as well quantified.

MORPHOLOGY OF SIC STOCKS The major SIC stocks are in petrocalcic and calcic horizons. It is estimated that calcic horizons contain 209 Pg C (1015 g C), petrocalcic horizons 104 Pg C, and calcareous soils not meeting the taxonomic criteria for a named carbonate horizon about 636 Pg C.[6] Disseminated carbonates (micritic or fine-grained carbonate particles) and segregated forms of soft masses, threads, films, or filaments are often present in soils at quantities insufficient to be recognized as calcic horizons.[8] Such disseminated and segregated carbonates may be the most chemically reactive form of SIC.

MAGNITUDE OF SIC STOCKS The amount of soil global inorganic carbon is large, with estimates ranging from 1576 to 695 Pg C.[3,15,16] Eswaran et al.[17] recently revised the estimates down to between 949 to 940 Pg C, which appear to be more reasonable (see Tables 1 and 2). Uncertainties in these numbers result because the depth of the soil represented has been arbitrarily set at 1 m, values represent combined lithogenic and pedogenic sources, and data for SIC stocks that are partitioned by soil taxa and global areal extent are approximate. Discrepancies in areal extents of soil orders=suborders, and the SIC stocks in Tables 1–4 are due to rounding errors and different sources of published information.

SIC STOCKS RELATED TO SOIL TAXONOMY ORDERS AND SUBORDERS Areal stocks of SIC in soil orders are presented in Table 1 and partitioned by the suborders of soil Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001706 Copyright # 2006 by Taylor & Francis. All rights reserved.

Inorganic Carbon: Global Stocks

891

Table 1 Soil inorganic carbon (SIC) and total soil carbon (TC) in soil orders of soil taxonomy TC Order

Pg

Alfisols

201

SIC % Global 8.1

Pg 43

% Global 4.5

SIC=TC % of order

SIC=TC % Global

21.4

1.7

Andisols

20

0.8

0

0.0

0.0

0.0

Aridisols

515

20.9

456

48.5

88.5

18.5

Entisols

353

14.3

263

28.0

74.5

10.6

Gelisols

323

13.1

7

0.8

2.2

0.3

Histosols

180

7.2

0

0.0

0.0

0.0

Inceptisols

224

9.1

34

3.6

15.2

1.4

Mollisols

237

9.5

116

12.3

48.9

4.7

Oxisols

126

5.1

0

0.0

0.0

0.0

64

2.6

0

0.0

0.0

0.0

Ultisols

137

5.6

0

0.0

0.0

0.0

Vertisols

64

2.6

21

2.3

32.8

0.9

0.0

Spodosols

Miscellaneous Total

24

1.0

0

0.0

2468

100.0

940

100.0

0.0 38.1


taxonomy in Table 2.[8] About 38% of the total carbon (TC) stocks are accounted for by SIC, with Aridisols, Entisols, and Mollisols contributing the most (Table 1). Nearly half of the SIC occurs in Aridisols (Table 1) with over 85% distributed among Argids, Calcids, and Cambids (Table 2). Entisols contain 28% of SIC (Table 1) with 77% associated with Orthents (Table 2). Mollisols contain the next largest amount of SIC ( 12%) with over 80% stored in Ustolls. The Alfisols contribute nearly 5% of the SIC stocks with over 95% stored in Ustalfs and Xeralfs. These distributions of SIC stocks reflect heavily on the semiarid and drier climates, coupled with carbonate-rich parent materials. Five of the soil orders contain no SIC stocks. These orders are the Andisols, Histosols, Oxisols, Spodosols, and Ultisols. They represent soils in subhumid to humid environments where carbonates have been leached into lower vadose zones, geologic substrata, and=or ground water aquifers. In these orders one would expect the parent materials to have been either noncalcareous or low in carbonate content. The ratios of SIC=TC as a percentage of the order and as a global percentage are given in Table 1. It is noteworthy that nearly 90% of the TC is SIC in Aridisols, 75% of the TC is SIC in Entisols, and about 50% of the TC is SIC in Mollisols. This reflects environments that do not favor sequestration of SOC because the soils are youthful, well oxidized, and have primary productivity constrained by long periods of soil moisture stress.


SIC STOCKS RELATED TO SOIL MOISTURE CONDITIONS Table 3 presents SIC and TC stocks according to global soil moisture and temperature conditions. Two soil moisture conditions account for 92% of the SIC stocks, namely arid ( 78%) and semiarid ( 14%). This is explained by the fact that these environments exhibit little soluble transport of pedogenic and lithogenic carbon stocks. In these environments, the proportion of TC that is SIC is 84% and 29%, respectively, indicating that accumulation of SOC is not as important in these dry climates as in humid and wetter regimes where SOC comprises 97% of the TC stock (Table 3). It is interesting to note that in Mediterranean climates with winter rain, SIC contributes 56% of the TC stock. Again, because of the semiarid condition the SOC pool is less important than SIC. From the above comments, it is not surprising that a little over 38% of the TC stock (2472 Pg C) is SIC (946 Pg C), with arid (30%) and semiarid (6%) regions serving as major contributors.

SIC STOCKS RELATED TO ECOLOGICAL REGIONS Table 4 illustrates SIC and TC stocks partitioned by ecological regions. Boreal and temperate regions contribute 82% of the global SIC, with tropical regions 16% and tundra regions 2%. The SIC is much higher

892


Table 2 Listing of soil taxonomy orders, suborders, ice-free land areas, and soil inorganic carbon (SIC) stocks Area ice-free landa Soil orders

Soil suborders

106 km2

%

Pg

%

Aqualfs Cryalfs Ustalfs Xeralfs Udalfs

12.620 0.836 2.518 5.664 0.896 2.706

9.65 0.64 1.92 4.33 0.69 2.07

43 0 1 29 12 1

4.6 0.0 0.1 3.1 1.3 0.1

0.91

0.70

0

0.0

Cryids Salids Gypsids Argids Calcids Cambids

15.728 0.943 0.890 0.682 5.407 4.872 2.931

12.02 0.72 0.68 0.52 4.13 3.73 2.24

456 17 21 23 135 165 95

48.5 1.8 2.2 2.4 14.4 17.6 10.1

Aquents Psamments Fluvents Orthents

21.137 0.116 4.428 2.860 13.733

16.16 0.09 3.39 2.18 10.50

263 0 33 27 203

28.0 0.0 3.5 2.9 21.6

Histels Turbels Orthels

11.26 1.01 6.33 3.91

8.60 0.77 4.84 2.99

7 4 3 1

0.8 0.4 0.3 0.1

1.53

1.17

0

0.0

Aquepts Cryepts Ustepts Xerepts Undepts

12.829 3.199 0.456 4.241 0.685 4.247

9.81 2.45 0.35 3.24 0.52 3.25

34 2

3.6 0.2

18

2.0

14

1.4

Albolls Aquolls Rendolls Xerolls Cryolls Ustolls Udolls

9.005 0.028 0.118 0.266 0.924 1.164 5.244 1.261

6.89 0.02 0.09 0.20 0.71 0.89 4.01 0.96

116 0 1 4 15 3 94 0

12.3 0.0 0.1 0.4 1.6 0.3 9.9 0.0

9.810

0.75

0

0.0

Alfisols

Andisols Aridisols

Entisols

Gelisols

Histosols Inceptisols

Mollisols

Oxisols Spodosols Ultisols Vertisols Aquerts Cryerts Xererts Torrerts Usterts Uderts Miscellaneous land Total a

(From Ref.[8].) (From Ref.[17].)

b


SICb

3.350

2.56

0

0.0

11.052

8.45

0

0.0

3.160 0.054 0.014 0.098 0.889 1.767 0.384

2.42 0.00 0.01 0.08 0.68 1.35 0.29

21 0 0 1 12 8 0

2.3 0.0 0.0 0.1 1.3 0.9 0.0

18.405

10.0

0

0.0

130.796

100.0

940

100.0


893

Table 3 Soil inorganic carbon (SIC) and total carbon (TC) stocks according to moisture conditions TC

SIC SIC=TC % Region

SIC=TC % Global

4.4

0.7

77.8

83.5

29.7

5.4

55.6

2.0

Moisture condition

Pg

% Global

Pg

Permafrost

405

16.4

18

1.9

Arid

877

35.5

732

90

3.6

50

Mediterranean

% Global

Semiarid

471

19.1

134

14.2

28.5

5.5

Humid

539

21.9

4

0.5

0.7

0.2

2.4

Perhumid Total

85

3.4

2

0.2

2472

100.0

946

100.0

0.1 38.2

(From Ref.[17].)

in the temperate region (55%) due to the larger extent of deserts. The small contributions of SIC to TC in tundra regions reflect cold temperatures favorable to SOC accumulation and dissolution of lithogenic SIC. Many of the tropical regions are more humid, and thus any SIC in the soil parent materials has been leached.

transport soluble products of carbonate weathering from the soil. Estimated results indicate that between 0.25 and nearly 0.40 Pg of SIC is lost per year globally in leaching calcareous environments. The major contributors to these leachates are the Entisol, Mollisol, and Alfisol soil orders. This value probably underestimates the actual flux of SIC stocks because it does not consider the amount of carbonic acid used to weather soil minerals or the amount that is flushed from soils of high leaching potentials, such as, Oxisols, Ultisols, and Spodosols, which occupy about 20% of the ice-free Earth’s surface.

FLUX OF SIC STOCKS An estimate of soluble carbonate fluxes can be made from mass balance reconstruction analysis of the soil carbonate phase. Two problems with this method are the lack of reliable estimates of soil chronology and assumptions of flux rates. The rates of soil genesis, and especially the more labile carbonate components, are known to be exponential functions, but the precise geometric rate function applicable is not known and would vary with different amounts of carbonate, carbonate mineralogy, leaching potential, soil temperature, soil permeability, and vegetative cover. The magnitude of leaching of soluble carbonates and bicarbonates has been estimated globally from calcareous soils.[12,13] These projections are based on a flux rate of between 7 and 8 g m2 yr1 from soil great groups likely to have leaching potentials sufficient to

CONCLUSIONS There is a very large global pool of SIC, but its formation, quantity, and dynamics are still uncertain. The major problems in determining SIC stocks are the incomplete global database and the difficulty in separating pedogenic from lithogenic carbonate sources. To provide better estimates of SIC stocks, more sampling and analysis are needed that involve good soil chronology, stable carbon isotopes to differentiate pedogenic from lithogenic carbonates, and techniques to quantify soluble SIC stocks.

Table 4 Soil inorganic carbon (SIC) and total carbon (TC) stocks according to ecological regions TC

SIC SIC=TC % Global

Pg

% Global

Tundra

405

16.4

18

1.9

4.4

0.7

Boreal

632

25.6

256

27.2

40.5

10.4

Temperate

873

35.4

518

55.1

59.3

21.0

Tropical

557

22.6

149

15.9

26.7

6.0

2466

100.0

940

100.0

Total (From Ref.[17].)


Pg

% Global

SIC=TC % Region

Ecological region

38.1

894

REFERENCES 1. Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Soil and Global Change; Advances in Soil Science; CRC Press, Inc.: Boca Raton, FL, 1995. 2. Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Soil Processes and C Cycles; Advances in Soil Science; CRC Press, Inc.: Boca Raton, FL, 1998. 3. Bouwman, A.F. Ed.; Soils and the Greenhouse Effect; John Wiley and Sons: New York, 1990. 4. Wisniewski, J., Sampson, R.N., Eds.; Terrestrial Biospheric Carbon Fluxes: Quantification of Sinks and Sources of CO2; 1993. 5. Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Soil Organic Matter in Temperate Agroecosystems; CRC Press: Boca Raton, FL, 1997. 6. Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Global Climate Change and Pedogenic Carbonates; Lewis Publishers: Boca Raton, FL, 2000. 7. Batjes, N.H. Total carbon and nitrogen in the soils of the world. European Journal of Soil Science 1996, 47, 151–163. 8. Soil Survey Staff. Soil Taxonomy: A Basic System of Soil Classification for Making and Interpreting Soil Surveys, 2nd Ed.; USDA-NRCS Agriculture Handbook 436; U.S. Government Printing Press: Washington, DC, 1999. 9. Nordt, L.C.; Hallmark, T.C.; Wilding, L.P.; Boutton, T.C. Quantifying pedogenic carbonate accumulations using stable carbon isotopes. Geoderma 1998, 82, 115–136. 10. Gile, L.H.; Grossman, R.B. The Desert Project Soil Monograph; USDA-SCS, National Technical Information Service: Springfield, VA, 1979.



11. Chadwick, O.A.; Capo, R.C. Partitioning Allogenic and Authigenic Sources of Carbonate in New Mexico Calcrete. Agronomy Abstracts, American Society of Agronomy, Cincinnati, Ohio, November 7–12, 1993; Madison, WI, 1993; 295 pp. 12. Nordt, L.C.; Wilding, L.P.; Drees, L.R. Pedogenic transformations in leaching soil systems: implications for the global carbon cycle. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, FL, 2000; 43–64. 13. Drees, L.R.; Wilding, L.P.; Nordt, L.C. Reconstruction of inorganic and organic carbon sequestration across broad geoclimatic regions. In Soil Carbon Sequestration and the Greenhouse Effect; Lal, R., Ed.; Soil Science Society of America Special Publication No 57; Madison, WI, 2001; 155–172. 14. Chadwick, O.A.; Kelly, E.F.; Merritts, D.M.; Amundson, R.G. Carbon dioxide consumption during soil development. Biogeochemistry 1994, 24, 115–127. 15. Eswaran, H.; Van den Berg, E.; Reich, P.; Kimble, J. Global soil carbon resources. In Soils and Global Change; Lal, R., Kimble, J., Levine, E., Steward, B.A., Eds.; CRC Press, Inc.: Boca Raton, FL, 1995; 27–44. 16. Sombroek, W.G.; Nachtergaele, F.O.; Hebel, A. Amounts, dynamics and sequestering of carbon in tropical and subtropical soils. Ambio. 1993, 22, 417–426. 17. Eswaran, H.; Reich, P.F.; Kimble, J.M.; Beinroth, F.H.; Padmanabhan, E.; Moncharoen, P. Global carbon stocks. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, FL, 2000; 15–26.

Inorganic Carbon: Land Use Impacts Donald L. Suarez Salinity Laboratory, United States Department of Agriculture-Agricultural Research Service (USDA-ARS), Riverside, California, U.S.A.

INTRODUCTION Land use affects soil inorganic carbon, but the changes are significant only over the long term. Both increases and decreases in carbon storage occur as a result of various management practices. In irrigated lands, the primary factors causing a decrease in inorganic carbon are high leaching and maintenance of elevated water content at or near the soil surface. These factors result in elevated carbon dioxide concentrations and thus increased dissolution of soil carbonates. Use of acidifying fertilizers such as ammonia and sulfur also act to reduce soil inorganic carbon. Practices that favor accumulation of carbonates in the soil include efficient irrigation with surface waters in arid and semiarid regions (leaching less than 30% of the applied water), irrigation with ground waters at elevated CO2 concentrations, application of gypsum to alkaline soils, and use of nitrate fertilizer. Other factors that affect soil carbonate content include land clearing, cropping practices, and erosion. Soil inorganic carbon constitutes a major carbon pool in the near surface environment. In arid regions, the inorganic carbon can comprise more than 90% of the total C in the soil. The major inorganic C mineral phases are calcite and dolomite. Both minerals are relatively insoluble, however, dolomite dissolution is much slower than calcite dissolution at the intermediate pH values relevant to soils. Also, dolomite does not readily precipitate under Earth surface conditions. As a result, dolomite content in soils will remain constant or decrease due to dissolution, while calcium carbonate content may either increase or decrease.

LAND CLEARING AND CROPPING Land clearing generally results in increased water runoff and soil erosion. This process or any other process such as tillage that increases erosion serves to remove the surface soil horizons. Since these horizons are generally depleted in inorganic carbon relative to less weathered, deeper horizons, there is an apparent increase in the inorganic carbon content of the surface soil as a result of erosion. In terms of carbonate dissolution the impact of land clearing is not certain. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001765 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

After clearing there is increased runoff, thus decreased infiltration, favoring less dissolution of carbonates. This effect may be compensated by the decreased water consumption (lower evapotranspiration) after clearing, resulting in increased deep recharge (greater carbonate dissolution). Depending on how much biomass remains after clearing, there is likely a short-term increase in soil CO2 followed by a longer-term reduction, favoring less carbonate dissolution in the soil. Land clearing and overgrazing in arid lands serve to increase wind erosion and to redistribute soil in the landscape. In this manner noncalcareous soils receive inputs of carbonates. This process increases net dissolution of carbonates in the landscape. In humid environments inorganic carbon is leached from the soil. The elevated CO2 concentrations in the soil enhance calcite solubility relative to Earth surface conditions. In humid environments carbonates are successively leached from the upper portions of the soil profile. Agricultural practices may serve to enhance or reduce the net removal of carbonates. Removal of vegetation from a site with practices such as tree harvesting or harvesting of forages serves to remove base cations and causes net acidification of the upper portions of the soil profile. If carbonates are present deeper in the soil, this acidification increases dissolution. The impact of removal of vegetation in humid environments with carbonates in the subsoil can be calculated by assuming that the net harvested alkalinity is compensated by an equal increase in carbonate dissolution in the subsurface.

FERTILIZATION Since optimum plant growth is generally at a pH lower than that observed in untreated calcareous soils, acid fertilizers are commonly applied. Use of sulfur with subsequent oxidation to sulfate results in acid release to the soil (2 mol of protons per mole of sulfur). Application of ammonia salts, with subsequent fixation into organic matter or oxidation to nitrous oxide or elemental nitrogen, also releases protons (2 and 1 mol of protons per mole of ammonia ions, respectively). This acidification will increase carbonate dissolution proportionately, and has a significant effect since ammonia salts are widely used. 895

896

Application of urea or ammonia gas should have no net effect on carbonate dissolution (upon oxidation to nitrous oxide or elemental nitrogen). In contrast, use of nitrate fertilizers serves to increase pH and thus reduce carbonate dissolution. Generally nitrate is not utilized on calcareous soils so the impact on inorganic carbon storage in soil is slight. The quantitative impact of fertilization on changes in inorganic carbon is not easily calculated, as it depends on the extent of N incorporation into organic matter,mineralization, the extent to which the harvested biomass is removed from the site, and the occurrence of carbonates in the subsurface. The addition of liming products, primarily calcite, are reported as 3.7 Tg C yr1 in the U.S. for 1978.[1] This is a significant but temporary addition to the soil carbon pool as it is assumed that the majority of the material is applied to acid soils and thus it is readily dissolved.

HUMID REGION IRRIGATION Irrigation in humid environments serves to increase the net recharge through the soils and thus should increase the removal of carbonates. These changes may be relatively difficult to detect in view of the limited amount of irrigation water added and the fact that irrigation in humid environments, although increasing rapidly, was very limited in the past. Field studies are needed to determine the impact of irrigation on changes in inorganic carbon storage in humid environments.

ARID REGION SOILS Arid zone soils usually contain at least minor amounts of carbonates, even if classified as noncalcareous. In the absence of irrigation, there may be redistribution of carbonates within the soil but little net precipitation. The majority of the pedogenic calcite is reprecipitated calcite with relatively small amounts added as a result of mineral weathering. Significant amounts of carbonates are also added to the surface of arid land soils as dust. Calcite is leached from the upper part of the soil profile by dissolution into the infiltrating rain, and is mostly reprecipitated at depth after plant extraction of the available water. Irrigation in arid and semiarid environments may result in a net increase or decrease in soil carbonate, depending on the water source and fraction of water applied that is leached (leaching fraction). There are two opposing effects. First, elevated CO2 concentrations in the root zone relative to the atmospheric condition results in enhanced calcite solubility and dissolution. Second, plant water extraction and evaporation concentrate the salts into a smaller volume of water and enhance calcite precipitation. At low leaching fractions,


Inorganic Carbon: Land Use Impacts

the effect of concentration of salts due to plant water extraction and evaporation is greater than the enhanced CO2 effect and there is net precipitation of calcite. For a calcite saturated surface water such as the Colorado River, it is estimated[2] that at a leaching fraction of 0.1 there is net precipitation of 125 kg ha1 yr1 of C, based on water consumption of 1.2 m yr1 and a CO2 partial pressure of 3 kPa. Model simulations indicated that net precipitation of calcite occurred as the sum of the loss of carbonates in the upper portion of the root zone and precipitation of calcite in the lower portion. At high leaching fractions there is net dissolution of carbonates. Using a calcite equilibrium model it is predicted that at a leaching fraction of 0.4 there will be a net dissolution of carbonates of 70 kg ha1 yr1 of C, again, based on water consumption of 1.2 m yr1 and CO2 partial pressure of 3 kPa.[3] In all instances there is a prediction of net dissolution in the upper portion of the soil profile and net precipitation in the lower portions of the profile. Using average leaching fractions for the western U.S., it is estimated that irrigation with surface waters on 12 million ha results in an increase in inorganic carbon of 1 Tg yr1,[4] or 80 kg ha1 yr1. Irrigation with groundwater saturated with respect to calcite will result in precipitation of carbonates at almost all leaching fractions, since the irrigation water is equilibrated at the groundwater CO2 partial pressure and is supersaturated upon degassing and application to the soil. Calcite saturated groundwater is used for irrigation on an estimated 3.12 million ha in the U.S. It is estimated that irrigation on these soils results in a net inorganic C precipitation of 1.3 Tg yr1 or 420 kg ha1 yr1.[3] The above calculations are dependent on several assumptions regarding calcite equilibrium and soil CO2 concentrations. Using a kinetic model for calcite precipitation and the measured CO2 partial pressure in the groundwater in Palo Verde Valley there is prediction of no net change in inorganic carbon at a leaching fraction of 0.5. Consistent with these predictions the groundwater composition in Palo Verde Valley shows no evidence for net precipitation or dissolution of carbonates in the soil. There is limited direct field evidence for the influence of irrigation on inorganic soil carbon and it is difficult to be certain that differences among sites are only related to changes in management. Researchers[4] observed a net decrease in the calcium carbonate content of three pairs of soil profiles taken from sites in Israel irrigated for approximately 40 yr as compared to nonirrigated sites. The estimated input of 4.40 m of water per year at those sites is contrasted with the yearly potential evapotranspiration of 1.93 m. The observed trend is qualitatively consistent with model predictions if we account for the input of rain and the estimated leaching fraction of 0.56. Isotopic evidence indicated that there was precipitation of

Inorganic Carbon: Land Use Impacts

pedogenic carbonate at depth despite a net decrease in carbonate content at depth. In a study in the San Joaquin Valley in California researchers compared samples of a soil taken from irrigated and native vegetation sites.[5] They also measured a net loss of carbonates attributed to 8 yr of irrigation. Net carbonate loss was estimated as 7 103 kg ha1 y1 (800 kg C ha1 yr1). Leaching fractions at the site were not reported but this value corresponds to approximately 10 times greater dissolution than expected based on model simulations. However, another study by the same author[6] found no change in total carbonate when comparing pedons with native vegetation and those irrigated for 5–25 yr. In this instance both gypsum and sulfur were applied as amendments for reclamation. Gypsum would tend to increase precipitation of carbonates while sulfur would acidify the soil and cause net dissolution of carbonates. In a recent, more extensive study[7] on paired soil cores (irrigated and adjacent nonirrigated sites) from the lower Colorado River basin there were no observed changes in net inorganic carbon storage after 90 yr of irrigation, and no isotopic C shifts indicative of recrystallization.

SODIC SOIL RECLAMATION Reclamation of sodic soils can result in either an increase or decrease in inorganic carbon in the soil. Gypsum application to a sodic and alkaline soil will increase the soil carbonate content, as the increased Ca will precipitate most of the soluble bicarbonate and carbonate. Application of sulfuric acid, sulfur, or green manuring all serve to dissolve soil carbonates. Green manuring as a reclamation practice consists of incorporating plant residues into the soil and leaching with water. The high CO2 production is combined with restricted gas transport creating very high CO2 concentrations in the soil, dissolving large amounts of calcium carbonate. It is estimated that this process can dissolve in the order of 400–800 kg ha1 during a year of reclamation. Use of acid is currently a widespread and generally recommended practice to prevent emitter clogging in drip irrigation systems. This practice may result in total removal of carbonates within 10–20 yr, for soils with less than 3% carbonate content.

IMPACT ON ATMOSPHERIC CARBON DIOXIDE Dissolution of carbonates in neutral to alkaline environments results in consumption of CO2 gas and formation of aqueous HCO3, while precipitation of carbonates results in release of CO2. The net effect of dissolution or precipitation of soil carbonate on atmospheric CO2 depends on the solution flow path.


897

In regions irrigated with surface water the dissolution of carbonates results in a net C sink. However, the high alkalinity drainage water usually flows back to the river. The resultant degassing of carbonic acid and reprecipitation of carbonate in the river or reservoir releases CO2 back into the atmosphere. If the water is recharged into deep aquifers the net soil flux is preserved. In acid environments, liming of soils results in CO2 release to the atmosphere as there is little or no net alkalinity produced.

CONCLUSIONS Land use practices have a long-term impact on soil inorganic carbon. Due to the large C pools in the soil these impacts are not generally observed in short-term studies. In humid environments the major anthropogenic impacts on inorganic C are liming of surface soils, use of NH4 vs. nitrate fertilizer, removal of vegetation, and erosion. In semiarid and arid environments increased inorganic C is favored by the use of groundwater for irrigation and application of gypsum. Decreased inorganic C is favored by inefficient irrigation with surface water and application of NH4 fertilizer. The net effect of irrigation on a global scale, neglecting the effects of fertilizer addition, is to increase soil inorganic C by 30 Tg C yr1 as well as to release an equal amount of C to the atmosphere. Liming practices in humid regions throughout the world are estimated to have no net effect on inorganic soil C but release up to 85 Tg C yr1 to the atmosphere.

REFERENCES 1. Voss, R.D. What constitutes an effective liming material. In National Conference on Agricultural Limestone; National Fertilizer Development Center: Muscle Shoals, AL, 1980; 52–61. 2. Suarez, D.; Rhoades, J. Effect of leaching fraction on river salinity. J. Irr. Drainage Div. ASCE 1977, 103 (2), 245–257. 3. Suarez, D. Impact of agriculture on CO2 as affected by changes in inorganic carbon. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J., Eswaran, H., Stewart, B., Eds.; Lewis: Boca Raton, FL, 1999; 257–272. 4. Magaritz, M.; Amiel, A. Influence of intensive cultivation and irrigation on soil properties in the Jordan valley, Israel: recrystallization of carbonate minerals. Soil Sci. Soc. Am. J. 1981, 45, 1201–1205. 5. Amundson, R.G.; Smith, V.S. Effects of irrigation on the chemical properties of a soil in the western San Joaquin valley, California. Arid Soil Res. Rehab. 1988, 2, 1–17. 6. Amundson, R.D.; Lund, L. The stable isotope chemistry of a native and irrigated typic natrargid in the San Joaquin valley of California. Soil Sci. Soc. Am. J. 1987, 51, 761–767. 7. Suarez, D.L. Impact of agriculture on soil inorganic carbon. Agron. Abstr., Madison, WI, 1998; 258–259.

Inorganic Carbon: Modeling Leslie D. McFadden University of New Mexico, Albuquerque, New Mexico, U.S.A.

Ronald G. Amundson University of California, Berkeley, California, U.S.A.

INTRODUCTION Virtually all carbon in soils of arid and semiarid regions of the world accumulates as pedogenic calcium carbonate (referred to subsequently as carbonate). The carbonate usually accumulates in layers that eventually attain the status of calcic horizons and, in much older soils, petrocalcic horizons. Numerous mechanisms for the accumulation of pedogenic carbonate in soils are recognized, but the most fundamental reason for accumulation is limited depth of soil-water movement and seasonally high evapotranspiration that favors precipitation of carbonate within the soil.[1] Many studies of calcic soils in the past few decades demonstrate a close correspondence between the depth of pedogenic carbonate accumulation and modern annual precipitation,[2–4] although recent studies show the relationship may be more complicated.[5] Other studies also show progressive, time-dependent accumulation in many environments;[1] these studies have led to the now generally accepted conceptual models of calcic soil development.[6] Development of a numerical model of carbonate accumulation, however, is a more challenging proposition, given the remarkably complex character of the soil system. Fortunately, certain aspects of calcic soils facilitate formulation of such numerical models. For example, the observed soil depth–climate relationship implies that, utilizing a sound strategy for simulation of water movement, determination of carbonate movement via solution transport is a reasonable proposition. In addition, a significant body of research shows that the majority of carbonate is derived from accumulated entrapped dust and Ca in rainwater.[1] Finally, data pertaining to calcite geochemistry and dissolution rates in different environments are available and show that, in soils associated with typical ranges in soil CO2, pH, and salinity, calcite is far more soluble than virtually all silicate minerals and has more rapid dissolution rates.[7] Consequently, a relatively simple model for carbonate movement in a soil based on relations in the CaCO3–H2O–CO2 system can be formulated that essentially ignores the more complex chemical reactions involving aluminosilicates. 898 Copyright © 2006 by Taylor & Francis

Research on the nature and composition of stable and unstable isotopes in pedogenic carbonate has also helped elucidate the nature of calcic soil development and improve the design for testing the results of numerical modeling.

THE COMPARTMENT MODEL AND SIMULATIONS OF PEDOGENIC CARBONATE ACCUMULATION The compartment-model, or ‘‘box-model,’’ approach to modeling of calcic soils accommodates continuously changing values among the interdependent variables that influence soil development. It enables integration of several factors that influence pedogenic carbonate accumulation and that can be explicitly considered in this model. These include soil–water movement and soil–water balance, changing soil CO2 concentrations and temperature with depth and season, initial parent material composition, carbonate and soluble salt additions from external sources, and calcite reactant surface area. The soil profile is represented by a vertical sequence of compartments of arbitrary dimensions, with the initial characteristics of each compartment specified (i.e., texture, available water-holding capacity, pCO2). A series of equations that enable forward modeling and simulation of evolving carbonate depth functions using the box-model approach can be derived on the basis of consideration of the factors indicated above. For example, the solubility of calcite is derived from the following equation, after Drever:[8] m3 Ca2þ ¼ ðpCO2 K1 Kcal KCO2 Þ= 4K2 gCa2þ þ g2 HCO3

ð1Þ

where Kcal is the calcite solubility product and K1, K2 and KCO2 are dissociation constants in the carbonate system, and gCa2þ and gHCO3 are the activity coefficients of Ca2þ and HCO3. A gravelly, permeable calcic soil probably best approximates an open system-weathering environment, in which case calcite dissolution rates are probably surface-area controlled Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001600 Copyright # 2006 by Taylor & Francis. All rights reserved.

Inorganic Carbon: Modeling

899

rather than diffusion controlled. Also, the dissolution rate is defined ultimately by the rate-limiting conversion of dissolved carbon dioxide (CO2 ) to H2CO3. At a very low solution volume to surface area ratios, and with fast, surface-controlled calcite dissolution rates, H2CO3 is rapidly depleted. In such circumstances, a commonly used rate equation that enables determination of dissolution rates in the CO2– CaCO3–H2O system[9] is: dC=dt ¼ ðA0 k=V Þð1 C=C Þn mg l1 s1

ð2Þ

where A0 is the surface area of rock in contact with water (cm2), V the water volume (cm3), k the reaction coefficient (mg cm l1 s1), n the reaction order, C the moles of calcite in solution, and C is the solubility of calcite. Values of n and k vary with saturation ratio, temperature, and pCO2. In the model, A0 =V can be specified depending on observed soil features. Eqs. (1) and (2) show that soil CO2 content is a very important variable, but soil CO2 contents may be highly variable.[10] Fortunately, studies show that a depth function for pCO2 that reflects prolonged seasonal respiration levels can be estimated, assuming the concentration of soil CO2 is described by mass transport of CO2 by gas diffusion.[11–12] The following diffusion-reaction equation, essentially Fick’s Second Law for a onedimensional case, is used in the model: @ Cs =@ t ¼ Ds @ 2 Cs =@ z2 þ fs ðzÞ ð3Þ where Cs is the concentration of CO2 in the soil (mol cm3), t the time (s), Ds the diffusion coefficient for CO2 in the soil (cm2 s1), z the depth in the soil (cm), and fs(z) is the production rate of CO2 as a function of depth (mol cm3 s1). At steady state, when @ Cs=@ t ¼ 0 ¼ Ds@ 2Cs=@ z2 þ fs, the general solution to this equation to produce a simple production function is: ð4Þ Cs ðzÞ ¼ f=Ds Lz z2 =2 þ C0 where C0 is the concentration of CO2 in the atmosphere (ppm) and L is the depth to the lower, no-flux boundary. Soil CO2 contents with depth calculated using this method are used to calculate carbonate solubility and dissolution rates with depth. Available water-holding capacity, infiltration, and percolation rates can be specified on the basis of laboratory soil measurements or estimated from field measurements or theoretical considerations. Earlier versions of the simulation model included certain assumptions that simplified numerical calculations, such as simple vertical saturated flow and constant soil temperature with depth. The lack of certain types of data (e.g., variation of pCO2 with depth and


time of year) also constituted a limitation on utility of the model. The model did enable simulation of 1) realistic depths and magnitudes of carbonate accumulation over thousands of years and 2) the range of effects of large climatic changes on calcic soils.[13–16] Model results emphasized the critical roles of external Ca2þ influx and influence of soil CO2 concentrations on carbonate accumulation. Model-simulated bimodal concentrations of carbonate based on theoretical, late-Pleistocene climatic conditions resembled those observed in late-Pleistocene, polygenetic soils; however, incompletely understood changes in the magnitude of climate changes, dust flux, and vegetation change in the Quaternary complicate attempts to simulate polygenetic soils.[7,14,16] Later versions of the model utilized important new inputs and employed routines that reflected improved understanding of key processes that strongly influence calcic soils. Studies of dust accumulation rates in the American Southwest,[17] C, O, and Sr isotopes in carbonate, and development of more sophisticated models for unsaturated flow in calcic soils[18] have allowed development of improved compartment models that can address new and more challenging research problems. For example, such numerical simulations demonstrate how climate changes in the Holocene might have dramatically influenced the rates and temporal patterns of soluble salt leaching and accumulation relative to pedogenic carbonate.[19] A more recent modeling study addressed the problem of how carbonate can occasionally accumulate at much shallower depths than those expected from the depth— annual leaching depth relationship.[20] This study showed how carbonate can be preferentially removed from depths of a few cm to a few dm below the soil surface, while carbonate simultaneously accumulates either as collars on surface pavement clasts or in the vesicular A horizon. Model results also explain how a significant change in climate or soil erosion rates could cause the dissolution of carbonate rinds on the tops and sides of boulders and=or the tops of limestone boulders at depths of up to several dm, unusual features observed in some calcic soils.[21]

ISOTOPES IN CALCIC SOILS During weathering, parent material carbonate undergoes dissolution and reprecipitation in the soil. The carbon (13C=12C, 14C=12C) and oxygen (18O=16O) isotope ratios of pedogenic carbonate that forms from dust or parent material carbonate, or from Ca2þ derived from silicate weathering, are determined by isotopic composition of soil CO2 and H2O. These are the primary carbon and oxygen reservoirs, respectively, for the carbonate. Therefore, pedogenic carbonate

900

reflects only isotopic conditions of the soil and bears no memory of the isotopic composition of the rock or mineral from which it was derived. Soil CO2 is derived primarily from decomposition of soil organic matter and root respiration. The C isotope composition of soil CO2 reflects: 1) the isotopic composition of these CO2 sources; 2) the effects of the diffusion of this CO2 toward the atmosphere; and 3) the isotopic composition of atmospheric CO2. In the 1980s, researchers recognized that fairly simple, steady state, diffusion models could be used to reasonably explain the observed depth patterns of C isotopes in both soil CO2 and pedogenic carbonate. The solution to the mathematical model that encompasses the forementioned processes describes the abundance of 12CO2 in soils. A related equation can be derived for 13CO2, and the ratio of the two models then describes the ratio of C isotopes at any given soil depth. A similar approach can also be used to model the 14C composition of CO2 with soil depth, with the additional complication that the two main sources of soil CO2 (humus decomposition and root respiration) have different 14C contents, making the solution to the model slightly more complex.[22] The C isotope diffusion model of soil CO2 has provided the opportunity to quantitatively use pedogenic carbonates in a number of applications: 1) paleovegetation studies;[23] 2) paleoatmospheric CO2 studies;[24] and 3) radiocarbon dating of pedogenic carbonate and geomorphic surfaces.[25–26] The O isotopic composition of soil water is determined by the O isotope composition of precipitation and the evaporation of soil water. It has been observed that the 18O content of modern precipitation is generally correlated with mean annual temperature on a global scale,[27] but regional differences due to storm sources can obscure these patterns.[28] If precipitation water (once stored in the soil) is subject to evaporation, an enrichment of the remaining soil water in 18O occurs because water vapor depleted in the ‘‘heavy’’ isotope is preferentially removed during evaporation. Models have been made that successfully explain the key components of this process[29] and the fact that soils subject to evaporation commonly have generally decreasing 18O contents of soil water with depth.[30] These models have two components. The first is a vapor transport layer (describing the flow of evaporating soil water to the atmosphere through a dry soil layer), and the second is an evaporating front layer. The evaporating front layer exists below the vapor transport zone. At the evaporating front 18O enrichment of soil water occurs as water is transferred to a vapor phase, and the remaining 18O-enriched soil water at the evaporating front then undergoes diffusional mixing with the less 18O-enriched water at greater depths. In general, these models have been more difficult to use than C isotope models due to the dynamic nature of soil



water (steady state assumptions are difficult to apply) and the array of model parameters, many of them not known with certainty for most soils. Oxygen isotopes in pedogenic carbonate have been used less extensively in paleoclimate work than C isotopes because of concern over possible evaporation of soil water that formed the carbonate. Amundson et al.[28] demonstrated that, except for hyperarid regions, the O isotope composition of pedogenic carbonate appears to reasonably reflect that of the local precipitation. There is a growing list of studies using carbonate O isotopes in Quaternary[28] and Tertiary[31] paleoclimate applications.

CONCLUSIONS Numerical models and isotope studies have proven to be valuable tools in the study of calcic soil development. They have helped elucidate the relation of climate, vegetation, and geomorphic processes to carbonate accumulation. The models are not able to explain the observed character of certain aspects of calcic soils, such as patterns of pedogenic carbonate development in some soil chronosequences[32], or the somewhat enigmatic formation of calcic soils in humid, monsoonal climates. Also, these models are not designed to simulate the evolution of very old, morphologically complex soils with petrocalcic horizons. Future models must be designed to address locally abundant calcic soils. Additional fieldwork and application of recently developed field and laboratory techniques will provide the basis for development of the next generation of numerical models.

REFERENCES 1. Birkeland, P.W. Soils and Geomorphology; Oxford University Press: New York, 1999; 720 pp. 2. Jenny, H. Factors in Soil Formation; McGraw-Hill: New York, 1941; 281 pp. 3. Arkley, R.J. Calculation of carbonate and water movement in soil from climatic data. Soil Sci. 1963, 96, 239–248. 4. Retallack, G.J. The environmental factor approach to the interpretation of paleosols. In Factors of Soil Formation: A Fiftieth Anniversary Retrospective; Amundson, R., Ed.; Special Publication 33; SSSA: Madison, WI, 1994; 31–64. 5. Royer, D.L. Depth to pedogenic carbonate horizon as a paleoprecipitation indicator? Geology 1999, 27, 1123–1126. 6. Machette, M.A. Calcic soils of the southwestern united states. In Soils and Quaternary Geology of the Southwestern United States; Weide, D.L., Faber, M.L., Eds.;


7.

8. 9. 10.

11.

12.

13.

14.

15.

16.

17.

18.

19.

Special Paper 203; Geological Society of America: Boulder, CO, 1985; 1–21. McFadden, L.D.; Tinsley, J.C. The rate and depth of accumulation of pedogenic carbonate accumulation in soils: formation and testing of a compartment model. In Soils and Quaternary Geology of the Southwestern United States; Weide, D.L., Faber, M.L., Eds.; Special Paper 203; Geological Society of America: Boulder, CO, 1985; 23–42. Drever, J.I. The Geochemistry of Natural Waters, 2nd Ed.; Prentice-Hall: Englewood Cliffs, NJ, 1991; 437 pp. Palmer, A.N. Origin and morphology of limestone caves. Geol. Soc. Am. Bull. 1991, 103, 1–21. Kiefer, R.H.; Amey, R.G. Concentrations and controls of soil carbon dioxide in sandy soil in the north Carolina coastal plain. Catena 1992, 19, 539–559. Solomon, D.K.; Cerling, T.E. The annual carbon dioxide cycle in a montane soil: observations, modeling, and implications for weathering. Water Resour. Res. 1987, 23, 2257–2265. Cerling, T.E.; Quade, J.; Yang, W.; Bowman, J.R. Carbon isotopes in soils and palaeosols, and ecology and palaeoecology indicators. Nature 1989, 341, 138–139. McFadden, L.D.; Amundson, R.G.; Chadwick, O.A. Numerical modeling, chemical, and isotopic studies of carbonate accumulation in soils of arid regions. Soil Sci. Soc. Am. Spec. Publ. 1991, 26, 17–35. McFadden, L.D. The impacts of temporal and spatial climatic changes on alluvial soils genesis in southern California. University of Arizona: Tucson; 430. Marion, G.M.; Schlesinger, W.H.; Fonteyn, P.J. Caldep: a regional model for soil CaCO3 (caliche) deposition in southwestern deserts. Soil Sci. 1985, 139, 468–481. Mayer, L.; McFadden, L.D.; Harden, J.W. Distribution of calcium carbonate in desert soils: a model. Geology 1988, 16, 303–306. Reheis, M.C.; Kihl, R. Dust deposition in southern Nevada and California, 1984–1989: relations to climate, source area, and source lithology. J. Geophys. Res. 1995, 100, 8893–8918. McDonald, E.V.; Pierson, F.B.; Flerchinger, G.N.; McFadden, L.D. Application of a soil–water balance model to evaluate the influence of holocene climatic change on calcic soils, Mojave desert, California, USA. Geoderma 1996, 74, 167–192. McFadden, L.D.; Crossey, L.J.; McDonald, E.V. Predicted response to calcic soil development to periods of significantly wetter climate during the late holocene (abstr.). Geol. Soc. Am. Abstr. Progr. 1990, 24 (6), A252.


901

20. McFadden, L.; McDonald, E.; Wells, S.; Anderson, K.; Quade, J.; Forman, S. The vesicular layer and carbonate collars of desert soils and pavements: formation, age and relation to climate change. Geomorphology 1998, 24, 101–145. 21. Treadwell-Steitz, C.; McFadden, L.D. Influence of parent material and grain size on carbonate coatings in gravelly soils, palo duro wash, New Mexico. Geoderma 2000, 94, 1–22. 22. Amundson, R.; Stern, L.; Baisden, T.; Wang, Y. The isotopic composition of soil CO2. Geoderma 1998, 82, 83–114. 23. Cerling, T.E. Development of grasslands and savannas in East Africa during the neogene. Palaeogeogr. Palaeoclimatol. Palaeoecol. 1992, 97, 241–247. 24. Cerling, T.E. Use of carbon isotopes in paleosols as an indicator of the P(CO2) of the paleoatmosphere. Global Biogeochem. Cycles 1992, 6, 307–314. 25. Amundson, R.; Wang, Y.; Chadwick, O.; Trumbore, S.; McFadden, L.D.; McDonald, E.; Wells, S.; DeNiro, M. Factors and processes governing the 14C content of carbonate in desert soils. Earth Planet. Sci. Lett. 1994, 125, 385–405. 26. Wang, Y.; McDonald, E.; Amundson, R.; McFadden, L.D.; Chadwick, O. An isotopic study of soil in chronological sequences of alluvial deposits, providence mountains, California. Geol. Soc. Am. Bull. 1996, 108, 379–391. 27. Rozanski, K.L.; Aragua´s-Aragua´s, L.; Gonfiantini, R. Isotopic patterns in modern global precipitation. In Climate Change in Continental Isotopic Records; Swart, P.K., Ed.; American Geophysical Union Monograph 78;: Washington, D.C., 1998; 1–36. 28. Amundson, R.; Chadwick, O.; Kendall, C.; Wang, Y.; DeNiro, M. Isotopic evidence for shifts in atmospheric circulation patterns during the late quaternary in midnorth America. Geology 1996, 24, 23–26. 29. Barnes, C.J.; Allison, G.B. The distribution of deuterium and 18O in dry soils. I. Theory. J. Hydrol. 1983, 60, 141–156. 30. Allison, G.B.; Hughes, M.W. The use of natural tracers as indicators of soil water movement in temperate semiarid regions. J. Hydrol. 1983, 60, 157–173. 31. Quade, J.; Cerling, T.E.; Bowman, J.R. Development of the asian monsoon revealed by marked ecological shift during the latest miocene in Northern Pakistan. Nature 1989, 343, 163–166. 32. Holliday, V.T.; McFadden, L.D.; Bettis, E.A.; Birkeland, P.W. Soil survey and soil geomorphology. In Profiles in the History of the U.S. Soil Survey; Helms, D., Efflard, A., Dwara, P., Eds.; Iowa State University Press: Ames, IO, 2001.

Insect Survival: Effect of Subzero Temperature Mike M. Ellsbury United States Department of Agriculture (USDA), Brookings, South Dakota, U.S.A.

INTRODUCTION Insects that inhabit soil in regions where winter temperatures fall below freezing have necessarily evolved behavioral and physiological mechanisms for surviving and developing in seasonally cold environments. Reviews of literature relating to the physiological, ecological, and behavioral adaptations that allow insects to survive cold temperature appear in Baust, Lee, and Ring,[1] Cannon and Block,[2] Block,[3] and Danks.[4] A recent review by Danks[5] considers relationships between dehydration and cold hardiness in dormant insects. Danks, Kukal, and Ring[6] and Block[7] have also compiled concise glossaries of scientific terms associated with cold hardiness in insects.

BEHAVIORAL ADAPTATION Behavioral mechanisms for avoidance of low temperatures may entail movement of freezing-susceptible overwintering stages deeper into the soil or to sites protected from cold temperatures.[8,9] Snow cover has an insulating effect that mediates winter soil temperatures, producing nearly isothermal soil temperature regimes when early snow cover occurs.[9,10] Where snow cover is sparse or lacking soil temperature more closely follows extremes of air temperature[11] and cold-hardiness adaptation becomes more important for overwinter survival of soil-dwelling insects. This is particularly true of soildwelling invertebrates of arctic regions that cannot avoid exposure to cold temperature and thus have necessarily evolved cold-hardiness traits and diapause capability that enable them to survive during exposure to subfreezing temperature. Where short warm seasons occur, modified, sometimes very prolonged, life cycles extending over more than one overwinter season are found in insects of cold environments.[3,6,9]

PHYSIOLOGICAL ADAPTATION Cold hardiness involves physiological and metabolic adaptations that may not always be concomitant with a depressed diapause metabolism. Diapause has been defined as hormonally controlled metabolic dormancy that occurs during a specific stage in the life cycle of an 902 Copyright © 2006 by Taylor & Francis

insect.[12–14] Thus, cold hardiness may occur independent of diapause, but often is an integral component of the diapause trait, and diapause expression may improve the cold-hardiness capability of the insect.[15]

FREEZE TOLERANCE VS. FREEZE SUSCEPTIBILITY Cold-hardy insects have been characterized into two groups: those that are tolerant to freezing and those that are susceptible to freezing.[8,16,17] The first of these, insects that are freeze tolerant, survive ice formation in body tissues but avoid damage to intracellular components because ice formation is usually confined to extracellular fluids. Protective mechanisms for freezing resistance include elevated solute levels, presence of nucleating agents, and accumulation of cryoprotectants in body fluids. The second broad category includes insects capable of undergoing supercooling without freezing of body fluids at temperatures that are below the true freezing point or melting point.[6] Insects that have the ability to supercool are considered freezing susceptible because they avoid freezing damage to body tissues only at temperatures above the supercooling point. Below the supercooling point, ice formation in body tissues results in death. These relationships are summarized in Fig. 1. It should, however, be noted that a strict dichotomous categorization of the cold-hardy stages of insects as either freeze tolerant or freeze susceptible may be an oversimplification.[7] Kostal and Havelka[18] observe that in soil environments where freezing events occur unpredictably, overwintering stages of some insects display both freeze susceptibility (supercooling) and freeze tolerance. In many insects, particularly those of more temperate environments, cold hardiness is associated with ability to survive low temperature during chilling events but is not necessarily associated with extremes of subfreezing temperature. The capacity of insects to supercool may be enhanced by acclimation at low temperatures, still above the supercooling point, prior to exposure to subfreezing temperatures. However, where acclimation does not occur, exposure to rapidly falling temperature, even for short durations, may induce cold shock or direct chilling injury. Cold shock is defined as stress associated with short duration or rapid exposure to low, but Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042706 Copyright # 2006 by Taylor & Francis. All rights reserved.

Insect Survival: Effect of Subzero Temperature

903

to survival in dry winter conditions. Block[26] concluded that resistance to desiccation may be considered preadaptive to cold hardiness in terrestrial arthropods.

CORN ROOTWORMS

Fig. 1 Generalized diagram showing insect responses to cold temperature. Heavy line indicates insect body temperature and shaded area is a zone where ice nucleation occurs. (From Ref.[6].)

nonfreezing temperature.[19] The rate of cooling is more significant in producing cold shock injury than are the duration or the level of low temperature attained during the chilling event. Insects subjected to cold shock sustain injury at temperatures well above the supercooling point. The mechanism of injury caused during cold shock probably involves changes in structure of lipids associated with membranes or thermoelastic stress induced by rapidly falling temperature.[19–21] Lee, Chen, and Denlinger,[22] described a process for rapid cold hardening of insects during exposure to low temperature that would otherwise result in cold shock. The phenomenon has been associated with glycerol accumulation[23] and production of stress proteins similar to those produced in response to heat shock.[24] The rapid cold hardening phenomenon was observed in nonoverwintering stages and probably provides insects with means to avoid direct chilling injury during rapid exposure to low temperatures above the supercooling point. This adaptation could be particularly important for soil-dwelling organisms of temperate regions that may be exposed to rapid freeze–thaw events during the autumn or spring. The association of heat-shock protein production with cold hardiness suggests similarities between rapid cold hardening and insect response to heat stress by production of heat-shock proteins. Mortality during cold exposure may result from dehydration during the freezing process, even in insects that are freeze tolerant.[8] Indeed, Ring and Danks[25] suggested that physiological and behavioral traits considered adaptive to cold hardiness may also be adaptive


The ability of certain Diabrotica beetle species (Coleoptera: Chrysomelidae) to survive winter desiccation in tropical climates may have preadapted them for survival in more northern climates. Pest Diabrotica, of the corn rootworm complex, are thought to have expanded their range northward from tropical regions into areas of more temperate climate concomitantly with the agricultural development of corn, Zea mays L., as a primary host plant. Western corn rootworms, Diabrotica virgifera virgifera LeConte, considered by Krysan[27] to be of tropical origin, survive cold winter conditions in temperate regions as a diapausing egg stage in the soil. The diapause mechanism that enabled survival of eggs during seasonal desiccation in the tropics also probably preadapted the soil-dwelling egg stage of this insect to survive in frozen soils.[28] A closely related subspecies, the Mexican corn rootworm, D. virgifera zeae, also overwinters as a diapausing egg stage in the soil but is adapted to survive dry winters in more southern climates.[29] These observations are consistent with the suggestion by Block[7] that insects tolerant of desiccation also are probably preadapted for survival of cold temperatures. In the northern Great Plains, the Diabrotica pest complex consists of two species, the western corn rootworm, D. virgifera virgifera, and the northern corn rootworm, D. barberi. Because both species overwinter as eggs in the soil, the occurrence of an economic infestation during the following growing season depends in part on the overwintering survival of the egg. Chiang,[30] Chiang, Mihm, and Windels,[31] Calkins and Kirk,[32] and Gustin[33,34] determined that winter soil temperatures reach levels low enough to cause mortality in overwintering western and northern corn rootworm eggs in the northern Great Plains. Nonetheless, a significant proportion of eggs laid each fall frequently survive winter conditions in corn-growing areas of North America to produce economic infestations of these insects each year. Ellsbury et al.[35,36] have shown that adult emergence patterns of both rootworm species have highly variable spatial distributions in the field. These distributions are probably the result of behavioral and mortality factors operating on eggs in the soil at two points in the seasonal life cycle of these insects. Firstly, ovipositional patterns may vary spatially in three dimensions with adult response to corn maturity, and soil texture and moisture conditions. Secondly, distributions of eggs in the soil may change from those that exist just

904

after the oviposition period because of differential overwinter mortality. The influences of fall tillage, soil characteristics, snow cover, and soil temperature are known to mediate egg mortality, but the interactions between these factors that produce spatial and temporal variation in surviving rootworm populations are poorly understood. The effect of low winter soil temperature on corn rootworm eggs has been the subject of intense study. Some early research on this subject was driven by the need for laboratory rearing of large quantities of postdiapause eggs for artificial infestation and thus was focused on the biology of egg diapause, temperature limits for development, and optimal chilling conditions for storage of eggs. Much research effort has been directed toward determining developmental thresholds and diapause requirements of rootworms in cold temperature environments; yet little information is available about cold hardiness and the actual cold hardening mechanism of the overwintering egg stage, nor is it known whether corn rootworm eggs can be considered freeze tolerant or freeze susceptible.

CONCLUSIONS Much of the current knowledge about how soil-dwelling insects survive cold soil temperature has come from scientific studies done in arctic or alpine environments under extreme conditions. The basic concepts arising from these studies have been summarized in this contribution. It is evident that insects have adapted to cold soil environments through a variety of mechanisms and that no single strategy suffices for survival in cold environments for all insects. The discussion of survival of corn rootworm eggs was intended to illustrate how some of this knowledge can be applied to an insect pest in a temperate agricultural setting. There are of course other soil-dwelling insects that overwinter in various life stages in both agricultural and natural settings that, for want of space, could not be considered here. However, it can be said that these basic concepts and references provided herein should serve as a point of departure for the reader wishing to undertake more in-depth study or research on the survival of insects in frozen soil environments.

REFERENCES 1. Baust, J.G.; Lee, R.E., Jr.; Ring, R.A. The physiology and biochemistry of low temperature tolerance in insects and other terrestrial arthropods: a bibliography. Cryo-Letters 1982, 3, 191–212.



2. Cannon, R.J.C.; Block, W. Cold tolerance of microarthropods. Biol. Rev. 1988, 63, 23–77. 3. Block, W. Cold tolerance of insects and other arthropods. Phil. Trans. R. Soc. Lond. B 1990, 326, 613–633. 4. Danks, H.V. Winter habitats and ecological adaptations for winter survival. In Insects at Low Temperature; Lee, R.E., Jr., Denlinger, D.L., Eds.; Chapman and Hall: New York, 1991; 31–259. 5. Danks, H.V. Dehydration in dormant insects. J. Insect Physiol. 2000, 46, 837–852. 6. Danks, H.V.; Kukal, O.; Ring, R.A. Insect cold-hardiness: insights from the Arctic. Arctic 1994, 47 (4), 391–404. 7. Block, W. Insects and freezing. Sci. Prog. 1995, 78 (4), 349–372. 8. Lee, R.E., Jr. Insect cold-hardiness: to freeze or not to freeze. Bioscience, 1989, 39 (5), 308–312. 9. Coulson, S.J.; Hodkinson, I.D.; Block, W.; Webb, N.R.; Worland, M.R. Low summer temperatures: a potential mortality factor for high arctic soil microarthropods. J. Insect Physiol. 1995, 41, 783–792. 10. Sharratt, B.S.; Baker, D.G.; Wall, D.B.; Skaggs, R.H.; Ruschy, D.L. Snow depth required for near steady state soil temperatures. Agric. Forest Meteorol. 1992, 57, 243–251. 11. Coulson, S.J.; Hodkinson, I.D.; Strathdee, A.T.; Block, W.; Webb, N.R.; Bale, J.S.; Worland, M.R. Thermal environments of arctic soil organisms during winter. Arct. Alp. Res. 1995, 27 (4), 364–370. 12. Beck, S.D. Insect Photoperiodism, 2nd Ed.; Academic Press: New York, 1980; 387 pp. 13. Saunders, D.S. Insect Clocks, 2nd Ed.; Pergamon Press: Oxford, 1982; 288 pp. 14. Tauber, M.J.; Tauber, C.A.; Masaki, S. Seasonal Adaptations of Insects; Oxford University Press: New York, 1986; 411 pp. 15. Denlinger, D.L. Relationship between cold hardiness and diapause. In Insects at Low Temperature; Lee, R.E., Jr., Denlinger, D.L., Eds.; Chapman and Hall: New York, 1991; 174–198. 16. Salt, R.W. Principles of insect cold-hardiness. Ann. Rev. Entomol. 1961, 6, 55–74. 17. Block, W. Cold hardiness in invertebrate poikilotherms. Comp. Biochem. Physiol. A 1982, 73, 581–593. 18. Kostal, V.; Havelka, J. Diapausing larvae of the midge Aphidoletes aphidimyza (Diptera: Cecidomyiidae) survive at subzero temperatures in a supercooled state but tolerate freezing if inoculated by external ice. Eur. J. Entomol. 2000, 97 (3), 433–436. 19. Denlinger, D.L.; Joplin, K.H.; Chen, C.-P.; Lee, R.E., Jr. Cold shock and heat shock. In Insects at Low Temperature; Lee, R.E., Jr., Denlinger, D.L., Eds.; Chapman and Hall: New York, 1991; 131–148. 20. Quinn, P.J. A lipid-phase separation model of low-temperature damage to biological membranes. Cryobiology 1985, 22, 128–146. 21. McGrath, J.J. Cold shock: thermoelastic stress in chilled biological membranes. In Network Thermodynamics, Heat and Mass Transfer in Biotechnology; Diller, K.R., Ed.; United Engineering Center: New York, 1987; 57–66.


22. Lee, R.E., Jr.; Chen, C.-P.; Denlinger, D.L. A rapid cold-hardening process in insects. Science 1987, 238, 1415–1417. 23. Chen, C.-P.; Denlinger, D.L.; Lee, R.E., Jr. Cold shock injury and rapid cold hardening in the flesh fly, Sarcophaga crassipalpis. Physiol. Zool. 1987, 60 (3), 297–304. 24. Burdon, R.H. Heat shock and the heat shock proteins. Biochem. J. 1986, 240, 313–324. 25. Ring, R.A.; Danks, H.V. Desiccation and cryoprotection: overlapping adaptations. Cryo-Letters 1994, 15, 181–190. 26. Block, W. Cold or drought—the lesser of two evils for terrestrial arthropods? Eur. J. Entomol. 1996, 93, 325–339. 27. Krysan, J.L. Diapause in the nearctic species of the virgifera group of Diabrotica: evidence for tropical origin and temperate adaptation. Ann. Entomol. Soc. Am. 1982, 75 (2), 136–142. 28. Krysan, J.L. Introduction: biology, distribution, and identification of pest Diabrotica. In Methods for the Study of Pest Diabrotica; Krysan, J.L., Miller, T.A., Eds.; Springer-Verlag: New York, 1986; 1–24. 29. Krysan, J.L.; Branson, T.F.; Castro, G.D. Diapause in Diabrotica virgifera (Coleoptera: Chrysomelidae): a comparison of eggs from temperate and subtropical climates. Ent. Exp. Appl. 1977, 22, 81–89. 30. Chiang, H.C. Survival of northern corn rootworm eggs through one and two winters. J. Econ. Entomol. 1965, 58 (3), 470–472.


905

31. Chiang, H.C.; Mihm, J.A.; Windels, M.B. Temperature effects on hatching of western and northern corn rootworms. Proc. North Cent. Br. Entomol. Soc. Am. 1972, 27, 127–131. 32. Calkins, C.O.; Kirk, V.M. Effect of winter precipitation and temperature on over-wintering eggs of northern and western corn rootworms. J. Econ. Entomol. 1969, 62 (3), 541–543. 33. Gustin, R.D. Soil temperature environment of overwintering western corn rootworm eggs. Environ. Entomol. 1981, 10 (4), 483–487. 34. Gustin, R.D. Diabrotica longicornis barberi (Coleoptera: Chrysomelidae): cold hardiness of eggs. Environ. Entomol. 1983, 12 (3), 633–634. 35. Ellsbury, M.M.; Woodson, W.D.; Clay, S.A.; Carlson, C.G. Spatial characterization of corn rootworm populations in continuous and rotated corn. In Precision Agriculture, Proceedings of the 3rd International Symposium on Precision Agriculture, Bloomington, Minnesota, Jun, 23–26, 1996; Robert, P.C., Rust, R.H., Larson, W.E., Eds.; American Society of Agronomy: Madison, Wisconsin, 1996; 487–494. 36. Ellsbury, M.M.; Woodson, W.D.; Clay, S.A.; Malo, D.; Schumacher, J.; Clay, D.E.; Carlson, C.G. Geostatistical characterization of the spatial distribution of adult corn rootworm (Coleoptera: Chrysomelidae) emergence. Environ. Entomol. 1998, 27 (4), 910–917.

Ion Exchange Bryon W. Bache Cambridge University, Cambridge, U.K.

INTRODUCTION Nature and Origin of Cation Exchange Many common substances exist as ions. The simplest example is common salt, sodium chloride (NaCl). This innocuous and life-sustaining compound is made of two dangerous and highly reactive elements, sodium and chlorine. When these two elements react together, the sodium atom loses an electron, giving it a positive charge (Naþ), while the chlorine atom gains an electron, giving it a negative charge (Cl). The positively charged ions are cations; the negative ones are anions. In a salt solution, the sodium and chloride ions are not physically bonded together but are free to diffuse independently in solution, attracted to each other only by electrostatic charges. Hence, the original derivation of the term ion (from the Greek for ‘‘wanderer’’) is now applied generally to any charged particle. Many substances in the soil solution exist as ions, but in a typical agricultural or garden soil most of the anions (the negatively charged particles) exist as large, insoluble macroions. Thus, the solid matrix of soil has a negative charge, and to maintain electrical neutrality, positively-charged cations are adsorbed onto the surface of the insoluble anions in a diffuseion swarm. This scenario facilitates cation exchange, whereby some of the adsorbed cations are exchanged for others.

CATION EXCHANGE Cation exchange is best illustrated by conducting a simple experiment. Place a layer of dry soil about 5 mm deep into a funnel, supported by a filter paper. Leach the soil slowly with demineralized water to remove the bulk of any soluble salts present, and discard this leachate. Then leach with 50 ml of water, but this time collect the leachate. (Although this will take some time, mixing the soil with coarse sand will speed up the process.) Next, leach the soil with 50 ml of a solution containing 2 to 3 g of ammonium chloride, and again collect the leachate. Test both the water leachate and the ammonium chloride leachate for calcium by adding ammonium oxalate solution. The water leachate will be free of calcium but the ammonium 916 Copyright © 2006 by Taylor & Francis

chloride leachate will give a white precipitate of calcium oxalate, showing that ammonium NH4þ has displaced calcium Ca2þ from the soil surfaces into solution. Ammonium has exchanged for calcium. A diagrammatic representation of this exchange reaction is shown in Fig. 1. Humus and clay are the two components of the soil matrix that contain the macroions responsible for cation exchange. Humus contains proton-donating functional groups on the surface of its molecules, which are undissociated at low pH but dissociate as the pH rises, to give a negatively charged surface. The most important of these is the carboxy group: RCOOH ! RCOO þ Hþ

ð1Þ

where R is the organic core. Clay consists of fine platy crystals of aluminosilicate minerals. The lattice of these minerals has a positive-charge deficiency caused by the isomorphous substitution of Al3þ for Si4þ, or Mg2þ for Al3þ in the structure of the mineral, which gives the crystal a net negative charge.

EXCHANGEABLE CATIONS, CATION EXCHANGE CAPACITY, AND BASE CATION SATURATION In most agricultural soils, the dominant exchangeable cation is calcium (Ca2þ), with lesser amounts of magnesium (Mg2þ), sodium (Naþ), and potassium (Kþ). These four are often termed base cations—i.e., the cations of the bases Ca(OH)2, Mg(OH)2, NaOH, and KOH. These should not to be confused with basic cations, of which MgOHþ is an example. In contrast, the acid cation aluminium Al3þ is dominant in acid mineral soils, and the hydrogen ion H3Oþ is dominant in acid organic soils. The total amount of cations that may be exchanged is the cation exchange capacity (CEC) of the soil. The sum of the positive charge on these cations equals the negative charge on the soil surface. The convention for CEC units most consistent with the International System of Units (SI) is ‘‘millimoles of charge per kilogram of dry soil’’ i.e., mmol(þ)kg1 for the cations and mmol()kg1 for the surface charge. A variety of experimental procedures has been developed to measure exchangeable cations and CEC.[1] CEC values Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001966 Copyright # 2006 by Taylor & Francis. All rights reserved.

Ion Exchange

Fig. 1 Diagrammatic representations of cation exchange: (A) the experiment described in the text; (B) the equilibrium between the surface and the solution.

in mmolkg1 may vary for clay minerals from 30 for kaolinite to 1200 for smectite, and for soils from 50 for Oxisols to 2000 for Histosols. The negative charge on humus is strongly pHdependent, so that apart from the nature and amounts of clay and humus, the most important variable affecting CEC is the pH of the soil (Fig. 2). Thus, when comparing different soils, CEC is conventionally measured at a standard pH, normally at the neutral point of pH 7.0 with an ammonium acetate buffer and ammonium NH4þ as the displacing cation, or at pH 8.2 with a triethanolamine buffer and barium Ba2þ as the displacing cation. Conventions for measuring CEC at a standard pH are appropriate for near-neutral soils dominated by permanent-charge clays. However, it is clear from Fig. 2 that they result in a highly inflated CEC value for acid soils when compared with the natural field soil, and they are even less appropriate for variable-charge soils, which acquire pH-dependent positive and negative charges (see Fig. 2 and Anion Exchange below). In these circumstances, CEC measured with unbuffered salt solutions at field pH provides a better interpretative value. This effective CEC (ECEC) equals the sum of the exchangeable cations extracted from the soil,


917

Fig. 2 The effect of pH on CEC: (a) a humic topsoil; (b) a mineral subsoil; (c) a strongly weathered Oxisol, showing both CEC and AEC.

including both base and acid cations. Thus, the pH of measurement should always be stated for CEC data. An important concept for acid soils is the base cation saturation (V). This is usually expressed as a percentage, and is then given by V ¼ 100 SMþ =CEC

ð2Þ

where SMþ is the sum of the exchangeable base cations, SMþ and CEC are measured in the same units, and CEC is at pH 7.0.

CATION EXCHANGE EQUILIBRIA AND CATION SELECTIVITY Cation exchange was first discovered by H. S. Thompson in 1850; a review of early related work was compiled by Kelley.[2] Good recent accounts are given by Mott,[3] Sposito,[4] and McBride.[5] Apart from the source of cation exchange (discussed earlier), most interest has focused on the relationship between the relative concentrations of different cations in solution and the amounts of surface-adsorbed cations, and how this varies for different exchange materials (e.g., silicate clay minerals, humus, and synthetic resins).

918

Ion Exchange

Fig. 1B illustrates the equilibrium between the adsorbed cations and those in the surrounding solution, all of which are free to diffuse within the constraints of electrical neutrality. This equilibrium is established rapidly, allowing the exchange reactions to occur when the solution concentrations are altered, as in the experiment described above, or when fertilizers, irrigation water, or environmental pollutants are added to soils. Theoretical approaches to quantify this equilibrium are important to predict cation behavior in soil processes. While rigorous thermodynamic treatments are complicated,[6] simplified empirical treatments are often more satisfactory. This author has found the most useful to be the corrected rational selectivity coefficient,[7] which incorporates the activities of the solution ions and the equivalent fraction of the adsorbed ions. Activity has been described as effective concentration and is usually a little lower than the actual concentration measured in mole=dm3. Equivalent fraction is the fraction of the total negative charge occupied by the adsorbed cation. In a simple homovalent exchange, such as potassium displacing sodium, soil Na þ Kþ

! soil K þ Naþ

ð3Þ

holes on the surface of silicate clays, thereby increasing the preference of the surface for potassium.

PRACTICAL IMPORTANCE OF CATION EXCHANGE This section outlines three of the many practical applications of cation exchange. Plant Nutrition[8] Many plant nutrients exist as cations: the four base cations mentioned above, ammonium NH4þ, and the heavy metal cations iron (Fe2þ), copper (Cu2þ), zinc (Zn2þ) and manganese (Mn2þ). Plants take up their nutrients as simple ions from solution, but when the solution is depleted, exchange desorption allows the adsorbed cations to provide a reserve from which the solution may be replenished (Fig. 1). To maintain electrical neutrality, desorption of one ion must be accompanied by adsorption of another, usually either Ca2þ (the most abundant cation) or the hydrogen ion H3Oþ, which plants excrete when they take up cation nutrients.

the selectivity coefficient K is given by Soil Development and Acidification[5,9]

XK ðNaþ Þ K ¼ XNa ðKþ Þ

ð4Þ

where X is the equivalent fraction of ion on the exchanger and the parentheses indicate the activity of that ion in solution. This selectivity coefficient shows the preference of the exchanger for one ion over another. In this example, if K is less than 1, sodium would be preferred by the exchanger, but in practice K is greater than 1 and potassium is preferred. For a heterovalent exchange reaction, such as calcium displacing sodium, soil Na2 þ Ca2þ

! soil Ca þ 2Naþ

ð5Þ

the selectivity coefficient is given by K ¼

XCa ðNaþÞ2 ; ½XNa 2 ðCa2þ Þ

ð6Þ

using the same conventions as previously. In general, cations with higher charge are more strongly adsorbed by exchangers (e.g., Al3þ > Ca2þ > Naþ) and for cations with similar charge, those with higher radius and lower hydration energy are preferred (e.g., Ba2þ > Ca2þ > Mg2þ). Structural considerations may also be involved. For example, the size of the potassium ion enables it to fit neatly into the


In a humid climate the course of soil development in a freely drained environment is always towards acidification. The hydrogen ion H3Oþ is strongly adsorbed and it exchanges for the base cations on the soil surfaces, which are then leached out, so that the soil gradually becomes more acid. The main sources of hydrogen ions are the dissolution of carbon dioxide, the release of organic acids when plant residues decompose, the deposition of atmospheric pollution, and the addition of fertilizer. Amelioration of Saline and Alkali Soils[5] An excess of sodium salts and=or exchangeable sodium in these soils causes problems. Amelioration is effected by adding gypsum (CaSO4) prior to leaching the soil with excess water. Calcium dissolves from the CaSO4 and exchanges for sodium (as in the example of heterovalent exchange 5, above), which is then leached in the drainage water.

ANION EXCHANGE Anion exchange[3–5] can occur on positively charged surfaces in a manner similar to cation exchange on

Ion Exchange

919

negative surfaces (e.g., as with synthetic resins based on quaternary ammonium groups.[7] Although there are no soil minerals that possess permanent positive charges and humus does not acquire a positive charge, strongly weathered acid soils containing amorphous hydrated iron and aluminum oxides can acquire a pH-dependent positive charge: FeOOH þ Hþ ! FeOþ þ H2 O

ð7Þ

This allows a small anion exchange capacity (AEC) to develop at a pH below 5, as illustrated in Fig. 2C, in addition to cation exchange on the negative surfaces of these variable-charge soils.[3,10] The main practical importance of anion exchange is that it allows the anions chloride Cl and nitrate NO3 to be resistant to leaching from the soil. This is particularly important for nitrate, which provides a reserve for plant nutrition. Phosphate has often been associated with anion exchange in the popular soil literature, but because its bonding to both neutral and positively charged sites is by hydroxyl ligand exchange rather than by diffuse-ion swarming, it is not genuinely ‘‘exchangeable’’ within the meaning of this article.


REFERENCES 1. Page, A.L.; Miller, R.H.; Keeney, D.R. Methods of Soil Analysis, Part 2, Chemical and Microbiological Properties; American Society of Agronomy: Madison, Wisconsin, 1982. 2. Kelley, W.P. Cation Exchange in Soils; American Chemical Society, Monograph No. 109; Reinhold: New York, 1948. 3. Mott, C.J.B. Surface chemistry of soil particles. In Russell’s Soil Conditions and Plant Growth, 11th Ed.; Wild, A., Ed.; Longman: Harlow, 1988; 239–281. 4. Sposito, G. The Chemistry of Soils; Oxford Univ. Press: New York, 1989. 5. McBride, M.B. Environmental Chemistry of Soils; Oxford Univ. Press: New York, 1994. 6. Sposito, G. The Thermodynamics of Soil Solutions; Clarendon Press: Oxford, 1981. 7. Helfferich, F. Ion Exchange; McGraw Hill: New York, 1962. 8. Mengel, K.; Kirkby, E.A. Principles of Plant Nutrition, 3rd Ed.; International Potash Institute: Bern, 1982. 9. Bache, B.W. Soil acidification and aluminium mobility. Soil Use and Management 1985, 1, 10–14. 10. Gillman, G.P.; Sumpter, E.A. Surface charge characteristics and lime requirement of soils derived from basaltic, granitic and metamorphic rocks in high-rainfall tropical queensland. Aust. J. Soil. Res. 1986, 24, 173–192.

Iron Oxides Nikolla P. Qafoku James E. Amonette Battelle–Pacific Northwest National Laboratory, Richland, Washington, U.S.A.

INTRODUCTION The iron (Fe) oxides are strongly pigmented clay-sized oxide, oxyhydroxide, and hydroxide minerals that occur in almost all soils and sediments. They are particularly abundant in highly weathered soils, where, in combination with Al and Ti oxides, they may account for the great majority of the soil mass[1] and the unique physical and chemical properties of these soils (Fig. 1). The sorptive properties of Fe oxides are largely determined by the interactions with the components of the soil solution such as protons and other dissolved ions. They also buffer the rich electron-transfer chemistry of Fe3þ and Fe2þ species. As a consequence of this sorptive and electrochemical reactivity, Fe oxides play a major role in the geochemical cycles of most elements having agronomic or environmental significance.

CLASSIFICATION AND OCCURRENCE The Fe oxides are compounds of Fe, O, and H that have structures based on close-packed arrays of O. The octahedral and tetrahedral cavities within these arrays are filled with either Fe3þ or Fe2þ to form Fe(O=OH)6, FeO6, or FeO4 structural units. The oxides are distinguished by the ways in which these structural units are arranged in space through corner, edge, or face sharing. Of the 15 Fe oxides currently recognized, only 8 are commonly found in soils and sediments.[2] Their selected properties are presented in Table 1.[1–3] Isomorphous Substitution All of the naturally occurring Fe oxide minerals will undergo some degree of substitution of other metal ions for Fe in their structures. The isomorphous substitution of Al3þ for Fe3þ is most frequent and especially common in goethite. However, other trivalent cations with similar ionic radii to Fe3þ (e.g., Ti3þ, Mn3þ, Co3þ, Cr3þ, and V3þ) may replace Fe3þ in the Fe oxide structures.[1] Cations with a higher (e.g., Ti4þ) or lower (e.g., Cu2þ, Zn2þ) valence may also 920 Copyright © 2006 by Taylor & Francis

substitute for Fe3þ but the extent of this substitution is not greater than 0.1 mol mol1.[4] Substitution of Mg2þ for Fe2þ also occurs in green rusts and magnetites.

METHODS OF IDENTIFICATION Bulk Analyses The presence and total quantity of Fe in a soil (i.e., FeT) can be determined by a variety of spectroscopic techniques following HCl or HF digestion, or nondestructively by energy-dispersive X-ray fluorescence (EDXRF). In some instances, determination of Fe valence is important, and this can be performed by wet-chemical techniques using phenanthroline or vanadate as indicators, or nondestructively using Mo¨ssbauer spectroscopy.[6] Relatively simple techniques may be used to identify Fe oxides in the field based on their typical colors and magnetic properties. In the laboratory, a variety of instrumental techniques can be used to confirm phase identity and to quantify amount. Of these, X-ray diffraction, infrared spectroscopy, electron microscopy, thermal analysis, and Mo¨ssbauer spectroscopy are the most commonly used techniques. Selective Extractions Operational measurements of Fe oxide content in soils are obtained using two selective-extraction techniques that discriminate between crystalline and poorly ordered oxides on the basis of solubility under different conditions. The dithionite-citrate-bicarbonate (DCB) method nominally dissolves all pedogenic Fe oxides (crystalline and poorly ordered). This Fe fraction is termed FeDCB. The ammonium oxalate extraction (conducted in the dark) essentially dissolves only the Fe fraction that is present in the poorly ordered Fe oxides (FeOX). The difference between FeDCB and FeOX is an estimate of the crystalline Fe oxide fraction (FeCRYS) present in soils. Highly weathered soils typically have high FeCRYS values. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006595 Copyright # 2006 by Taylor & Francis. All rights reserved.

Iron Oxides

921

Fig. 1 Fe-rich highly weathered soils from the southeastern United States. [(A) Tifton serie (fine-loamy, siliceous, thermic Plinthic Kandiudults); (B) Faceville serie (clayey, kaolinitic, thermic Typic Kandiudults)]. (View this art in color at www.dekker.com.)

The Fe in silicate minerals (FeSIL), which is not dissolved appreciably by the DCB method, can be calculated as the difference of FeT and FeDCB. This Fe source may be mobilized during weathering and the ratio of FeSIL to FeT is a useful parameter in estimating the stage of soil weathering. Highly weathered soils (e.g., Oxisols and Ultisols) have small FeSIL=FeT values.

FORMATION AND CHEMISTRY Formation The process of Fe oxide formation in soils (Fig. 2) starts with the oxidation of structural Fe2þ or aqueous Fe2þ that is released during the weathering of Fe2þbearing primary minerals. In addition, Fe3þ may be released upon dissolution of Fe3þ oxides and Fe3þbearing minerals. Ferromagnesian silicates (e.g., olivines, pyroxenes, amphiboles, and micas) and Fe– Ti oxides are the primary Fe-bearing minerals and


serve as a weatherable source of Fe in soils. Usually, minerals that contain Fe2þ are less stable because structural Fe2þ may oxidize to Fe3þ, creating a charge imbalance that weakens the structure of the mineral and promotes dissolution.[2] Depending on the rate and extent of oxidation, the pH, and the presence of various anions in the solution, ferrihydrite, goethite, schwertmannite, or green rust will directly precipitate from the solution. Rapid oxidation, which results in a higher Fe3þ activity in the soil solution, promotes the formation of ferrihydrite. Goethite forms when the soil is rich in organic matter, inorganic colloids, or biological exudates that may sorb or complex Fe3þ, thus decreasing the activity of the free aqueous ions. Schwertmannite forms at low pH and high SO42 typical of acid mine drainage. Green rust forms when oxidation is very slow and incomplete, and when pH is neutral or higher. Continued slow oxidation of green rust under high CO32 conditions and dissolution of ferrihydrite and schwertmannite eventually lead to formation of goethite. Rapid oxidation of green rust under slightly acidic, low-CO32 conditions

922

Iron Oxides

Table 1 Selected properties of iron oxides commonly found in soils and sediments Chemical formula

Minerals Oxides

Oxyhydroxides

Hydroxides

Hematite

a-Fe2O3

Magnetite

Fe3O4

Maghemite

g-Fe2O3

Goethite

a-FeOOH

Lepidocrocite

g-FeOOH

Schwertmannite

Fe8O8(OH)6SO4

Green Rust

variable (Fe62þFe23þ) (OH)16CO3 2H2O

Ferrihydrite

Fe5HO8 4H2O

Occurrence/ formation

Other properties

Humid tropical, subtropical and Mediterranean soils (low organic matter and warm temperatures) Inherited from parent materials, but also produced by bacterial activity Highly weathered soils, especially Oxisols; formed by the oxidation of magnetite or by heating of goethite in the presence of organic matter

Blood-red, fine-grained, hexagonal plates

High stability and crystallinity

Black, coarsely grained cubes

Contains both Fe2þ and Fe3þ, ferrimagnetic

Brown, coarsely grained cubes

Has magnetite structure, ferromagnetic, minor Fe2þ

Most common Fe oxide in soils, especially under cool and wet conditions, rich in organic matter Seasonally anaerobic, noncalcareous soils in cool areas Acid-sulfate soils, and waters draining mine soils and pyritic rocks Transitory phase formed on rapid mixing of Fe2þ and Fe3þ salts under alkaline to neutral, slightly oxic conditions; may be produced by bacterial activity Organic-matter-rich soils subject to a rapid change from anoxic to oxic conditions; imprecise chemical formula

Yellow, fine-grained stubby crystals with little acicularity

Highly stable, Al substitution common

Distinctive orange-colored laths

Similar to goethite, but little or no Al

Dark-yellow–orange clusters of fibrous crystals

Sulfate anion occupies tunnel structure

Distinctive blue-green, very fine-grained hexagonal plates

Contains Fe2þ and Fe3þ, reacts readily with oxidants; may contain Mg and anions besides CO32

Brownish-orange spheres, 2–3 nm in diameter

Very poorly ordered, high surface area

forms lepidocrocite. Slow oxidation of green rust under slightly alkaline, low-CO32 conditions forms magnetite. Hematite forms by a solid state transformation of ferrihydrite (predominantly) or magnetite.


Color and morphology

Maghemite forms by the oxidation of lithogenic magnetite or by the transformation of other Fe oxides by heating above 300 C in the presence of organic compounds. Hematite, maghemite, lepidocrocite, and

Iron Oxides

923

Fig. 2 Fe oxides formation and transformation in soils. (From Ref.[2].) (View this art in color at www.dekker.com.)

goethite are the most stable Fe oxides. The abundance of goethite in soils stems from the many pathways by which it can form, coupled with its high thermodynamic stability.[2]

Chemistry Fe oxides are one of the end products of the weathering process in soils. They form clay-sized particles as small as a few nanometers across (e.g., ferrihydrite) and have specific surfaces ranging as high as several hundreds of m2 g1. Even crystalline Fe oxides (e.g., goethite) may have specific surfaces of several tens of m2 g1. As oxides, the functional groups on their surfaces may have positive, negative, or no charge depending on the pH and on the concentration and nature of other ions in the contact solution. A net positive surface charge is usually observed in soils because Fe oxides have a point-of-zero-charge in the neutral or slightly basic pHs. The functional groups on the


surface form complexes with cations and anions from the aqueous phase. Oxyanions (e.g., phosphate) are particularly strongly bound as they form innersphere complexes in which an anion oxygen becomes part of the oxide surface structure. The positive charge on the oxide surfaces may also be electrostatically attracted to the permanent negative charge on layered aluminosilicate clay minerals. Formation of electrostatic bonds between oxide and phyllosilicate minerals promotes the formation of soil aggregates and stabilizes them, especially in highly weathered variable-charge soils (Fig. 3). Fe oxides also buffer the oxidation–reduction activity of soils. The reduction potential for the aqueous Fe3þ=Fe2þ couple (þ0.77 V) is situated in the middle of the limits for aqueous chemistry (i.e., 1.1 to þ1.8 V), and Fe3þ in the poorly ordered oxides serves as a terminal electron acceptor for microorganisms once oxygen and nitrate are consumed. Under these anoxic conditions, large quantities of soluble Fe2þ may be generated and become mobile. The crystalline

924

Iron Oxides

Fig. 3 Interactions among positively charged Fe oxides and negatively charged phyllosilicates in Fe-rich soils. (View this art in color at www.dekker.com.)

Fe oxides are less susceptible to reductive dissolution and can retain some of these mobile Fe2þ cations on their surfaces, where they become much more effective reductants for environmental contaminants.[7–9] The concentrations of aqueous Fe2þ in soils may be as high as 104 mol L1, whereas that of Fe3þ is typically much lower (about 1012 mol L1) because of the low solubility of Fe3þ oxides at neutral and alkaline pH.

FURTHER READING For a more detailed treatment of the classification, structure, and chemistry of the iron oxides in soils, several excellent reviews in Refs.[1–3] and in the two books, Refs.[4,5] are recommended. For discussion of the methods used to characterize Fe oxide and other soil minerals, see Ref.[6]. Detailed discussions of the environmental chemistry of Fe are found in Refs.[7–10].


CONCLUSIONS Fe oxides are one of the most important mineralogical components in soils. They primarily form by the oxidation of Fe2þ that is dissolved during the weathering processes. Because they occur as clay-sized particles having high specific surfaces, they substantially contribute to the physical and chemical properties of the soil. Their sorption and electron-buffering properties significantly affect the geochemical cycles of almost all elements having agronomic or environmental significance.

REFERENCES 1. Ka¨mpf, N.; Scheinost, A.C.; Schulze, D.G. Oxide minerals. In Handbook of Soil Science; CRC Press: Boca Raton, FL, 2000; F-125–F-137. 2. Bigham, J.M.; Fitzpatrick, R.W.; Schulze, D.G. Iron oxides. In Soil Mineralogy with Environmental

Iron Oxides

Applications; Dixon, J.B., Schulze, D.G., Eds.; SSSA Book Series; Soil Science Society of America, Inc.: Madison, WI, 2002; Vol. 7, 323–366. 3. Schwertmann, U.; Taylor, R.M. Iron oxides. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; SSSA Book Series; Soil Science Society of America, Inc.: Madison, WI, 1989; Vol. 1, 379–438. 4. Cornell, R.M.; Schwertmann, U. The Iron Oxides; VCH Publ.: Weinheim, Germany, 1996. 5. Stucki, J.W.; Goodman, B.A.; Schwertmann, U., Eds.; Iron in Soils and Clay Minerals; D. Reidel Publ. Co.: Dordrecht, Holland, 1988.


925

6. Amonette, J.E.; Zelazny, L.W., Eds.; Quantitative Methods in Soil Mineralogy; Miscellaneous Publ., Soil Science Society of America: Madison, WI, 1994. 7. Amonette, J.E. Iron redox of clays and oxides: Environmental applications. In Electrochemical Properties of Clays; Fitch, A., Ed.; CMS Workshop Lectures; Clay Minerals Society: Aurora, CO; Vol. 10, 89–146, in press. 8. Stumm, W.; Morgan, J.J. Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters, 3rd Ed.; Wiley: New York, 1996. 9. Stumm, W. Chemistry of the Solid–Water Interface; Wiley: New York, 1992. 10. Sposito, G. The Surface Chemistry of Soils; Oxford University Press: New York, 1984.

Irrigation Jack Keller Utah State University, Logan, Utah, U.S.A.

INTRODUCTION The basic idea behind irrigation is to supply water to and store water in soils for later use by plants where or when insufficient water for the desired level of plant growth is supplied by natural precipitation. There are three basic methods used for supplying irrigation water to the surface of soils and two for applying water beneath the soil surface. IRRIGATION METHODS The three application methods for applying water to the surface of soils are: 1) surface irrigation, where the soilsurface is used to convey the irrigation water to where it will be stored in the soil; 2) sprinkle irrigation, where the water is sprinkled on the soil surface where it is to be stored; and 3) drip irrigation, where the water is directly applied by dripping or being sprayed from small holes in the pipe network to where it is to be stored in the soil. The two methods for applying irrigation water beneath the soil surface are: 1) subsurface drip irrigation, where the network of pipe and outlets are buried beneath the soil surface and 2) subirrigation, where the water table is controlled and the water reaches the plant root system through capillary action. These methods of irrigation and irrigation efficiency issues are discussed in detail in other sections of this encyclopedia. WATER STORAGE CAPACITY OF SOILS Understanding basic soil–water–plant relations is central to the ability to design and manage irrigation systems. The tables related to soil water presented in this section provide useful guidelines for estimating irrigation requirements and for preliminary irrigation design purposes. However, more site-specific information should be used for designing and managing irrigation systems and evaluating irrigation system performance.

economic loss. Table 1 gives typical ranges of available water holding capacities (field capacity minus permanent wilting point) of soils of different textures. Table 1 should only be used as a guide for preliminary system designs and management purposes. Final designs and management decisions should be based on actual field data. Root Depth The total amount of soil water available for plant use in any soil is the sum of the available water-holding capacities of all soil horizons occupied by plant roots. Typical plant feeder root and total root depth are given in many references; however, the actual depths of rooting of the various crops are affected by soil conditions. Therefore, the root depth for any set of crop and site should be checked. Where local data are not available and there are no expected root restrictions, Table 2 can be used as a guide to estimating the effective root depths. The values given are taken from Ref.[2] and are based on the author’s estimate of averages selected from a large number of references. They represent the depth at which crops will obtain the major portion of the water they need when grown in a deep, well-drained soil that is adequately irrigated. CONSUMPTIVE USE AND DESIGN Deciding how much water an irrigation system should be able to deliver to a crop over a given period is ultimately a question of selecting a capacity that will maximize profits to the farmer. To begin to address this question of system capacity, it is necessary to know how much water a crop will use over the entire growing season and during the part of the season when water use is at its peak. The rate of water use during this peak period provides the basis for determining the rate at which irrigation water should be delivered to the field. Examples of typical seasonal and peak daily crop water requirements are given in Table 3.

Soil Water Holding Capacity

SOIL MOISTURE MANAGEMENT

Soils of various textures have varying abilities to retain water. Except for required periodic leaching, any irrigation beyond the field capacity of the soil is an

A general rule of thumb for many field crops in arid and semiarid regions is to maintain the soil moisture deficit (SMD) within the root zone above 50%


Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001577 Copyright # 2006 by Taylor & Francis. All rights reserved.

Irrigation

927

Table 1 Range in available water-holding capacity of soils of different texture Water-holding capacity

Irrigation Depth The maximum net depth of water to be applied per irrigation, dx, is the same as the maximum allowable depletion of soil water between irrigations. Assuming no additional water is necessary for leaching purposes, it is computed by:

Range (mm/rn)

Average (mm/rn)

Very coarse texture— very coarse sands

33–62

42

Coarse texture—coarse sands, fine sands, and loamy soils

62–104

83

dx ¼

Moderately coarse texture—sandy loams

104–145

25

Medium texture—very fine sandy loams, loams, and silt loams

125–192

167

Moderately fine texture— clay loams, silty clay loams, and sandy clay loams

145–208

183

Fine texture—sandy clays, silty clays, and clays

133–208

192

where dx Maximum net depth of water to be applied and stored in the root zone per irrigation event [mm (in.)]. MAD Management-allowed deficit, which can be estimated from Table 4 (%). Wa Available water-holding capacity of the soil, which can be estimated from Table 1 [mm=m (in.=ft)]. Z Effective root depth, which can be taken from Table 2 [mm (ft)].

Peats and mucks

167–250

208

Irrigation Interval

Soil texture

MAD Wa Z 100

ð1Þ

[1]

(Adapted from Ref. .)

The appropriate irrigation interval, which is the time that should elapse between the beginning of two successive irrigation events, is determined by: of the total available water-holding capacity. This is a management-allowed deficit (MAD ¼ 50%) because it is also desirable to bring the moisture level back to field capacity with each irrigation. (In humid regions, it is necessary to allow for rains during the irrigation period. However, the 50% limitation on soil moisture depletion should be followed as a general guide for field crops.) Soil management, water management, and economic considerations determine the amount of water used in irrigating and the rate of water application. The standard approach has been to determine the amount of water needed to fill the entire root zone to field capacity and then to apply a sufficiently larger amount to account for evaporation, leaching, and efficiency of application. The traditional approach to the frequency of application has been to take the depth of water in the root zone reservoir that can be extracted assuming MAD ¼ 50% and, using the daily consumptive use rate of the crop, to determine how long this supply will last. Such an approach is useful only as a guide to irrigation requirements, as many factors affect the volume and timing of applications for optimal design and operation of a system. Table 4 is presented as a guide for selecting the appropriate MAD or for near optimum production of various crops. As indicated in Table 4 for crops with a high market value, it is often profitable to irrigate well before half of the Soil moisture in the root zone has been depleted, i.e., SMD ¼ 50%.


f0 ¼

dn Ud

ð2Þ

where f 0 ¼ Irrigation interval or frequency (days). dn ¼ Net depth of water applied and stored in the soil per irrigation event, to meet consumptive use requirements [mm (in.)]. Ud ¼ Conventionally computed average daily crop water requirement, or use rate, during the peak-use month, which can be estimated from Table 3 [mm=day (in.=day)]. The value selected for dn will depend upon system design and environmental factors and should be equal to or less than dx. When dn is replaced by dx in Eq. (2), f 0 becomes the maximum irrigation interval, fx. IRRIGATION SYSTEM DESIGN, MANAGEMENT AND SCHEDULING General Design Concepts Design of all irrigation systems is process of synthesis, where properties, such as a soil’s intake rate and crop water requirements; items, such as canals, pipes, and pumps; and processes, such as trenching, coaxing water down furrows, or moving pipe, must be integrated to form a good irrigation system. The irrigation

928

Irrigation

Table 2 Effective crop root depths that would contain approximately 80% of the feeder roots in a deep, uniform, well-drained soil profilea

Table 2 Effective crop root depths that would contain approximately 80% of the feeder roots in a deep, uniform, well-drained soil profilea (Continued)

Crop

Root depth (M)

Crop

Root depth (M)

Alfalfa

1.2–1.8

Peach

0.6–1.2

Almonds

0.6–1.2

Peanuts

0.4–0.8

Apple

0.8–1.2

Pear

0.6–1.2

Apricot

0.6–1.4

Pepper

0.6–0.9

Artichoke

0.6–0.9

Plum

0.8–1.2

Asparagus

1.2–1.8

Potato (Irish)

0.6–0.9

Avocado

0.6–0.9

Potato (sweet)

0.6–0.9

Banana

0.3–0.6

Pumpkin

0.9–1.2

Barley

0.9–1.1

Radish

Bean (dry)

0.6–1.2

Safflower

0.9–1.5

Bean (green)

0.5–0.9

Sorghum (grain & sweet)

0.6–0.9

Bean (lima)

0.6–1.2

Sorghum (silage)

0.9–1.2

Beet (sugar)

0.6–1.2

Soybean

0.6–0.9

Beet (table)

0.4–0.6

Spinach

0.4–0.6

Berries

0.6–1.2

Squash

0.6–0.9

0.3

Broccoli

0.6

Strawberry

0.3–0.5

Brussel sprouts

0.6

Sudan grass

0.9–1.2

Cabbage

0.6

Tobacco

0.6–1.2

Cantaloupe

0.6–1.2

Tomato

0.6–1.2

Carrot

0.4–0.6

Turnip (white)

0.5–0.8

Cauliflower

0.6

Walnuts

1.7–2.4

Celery

0.6

Watermelon

0.6–0.9

Chard

0.6–0.9

Wheat

0.8–1.1

Cherry

0.8–1.2

Citrus

0.9–1.5

Coffee

0.9–1.5

Corn (grain & silage)

0.6–1.2

Corn (sweet)

0.4–0.6

Cotton

0.6–l.8

Cucumber

0.4–0.6

Eggplant

0.8

Fig

0.9

Flax

0.6–0.9

Grapes

0.5–1.2

Lettuce

0.2–0.5

Lucerne

1.2–1.8

Oats

0.6–1.1

Olives

0.9–1.5

Onion

0.3–0.6

Parsnip

0.6–0.9

Passion fruit

0.3–0.5

Pastures

0.3–0.8

Pea

0.4–0.8 (Continued)


a

Approximately 80% of the feeder roots are in the top 60% of the soil profile. Soil and plant environmental factors often offset normal root development; therefore, soil density, pore shapes and sizes, soil–water status, aeration, nutrition, texture and structure modification, soluble salts, and plant root damage by organisms should all be taken into account. (Adapted from Ref.[2].)

designer’s art is to know the kinds of hardware and management techniques appropriate for a given cropping system and site and to have a clear mental image of what the system can accomplish and how the completed system will appear. Careful site analysis is essential as it provides data that lead to an understanding of the physical, economic, and social resources that determine what can and ought to be accomplished by a proposed irrigation system. Generally irrigation systems are designed to meet average peak-water-use requirements. Sometimes, to reduce costs or to stretch limited water supplies, systems are designed to optimize production per unit of water applied. In such cases systems can be designed to apply only about 80% of peak water requirements and still obtain up to 95% of optimum yields. For deep-rooted crops in fine-textured soils, an appreciable

Irrigation

929

Table 3 Typical peak daily and seasonal crop water requirements in different climates Type of climate and water requirements (mm) Cool Crop

Moderate

Hot

High desert

Low desert

Daily

Seas

Daily

Seas

Daily

Seas

Daily

Seas

Daily

Seas

Alfalfa

5.1

635

6.4

762

7.6

914

8.9

1016

10.2

1219

Pasture

4.6

508

5.6

610

6.6

711

7.6

762

8.9

914 a

508a

Grain

3.8

381

5.1

457

5.8

508

6.6

533

5.8

Beets

4.6

584

5.8

635

6.9

711

8.1

732

9.1

914

Beans

4.6

330

5.1

381

6.1

457

7.1

508

7.6

559

Corn

5.1

508

6.4

559

7.6

610

8.9

660

10.2

762

10.2

Cotton

—

—

6.4

559

7.6

660

—

—

Peas

4.6

305

4.8

330

5.1

356

5.6

356

5.1b

Tomatoes

4.6

457

5.1

508

5.6

559

6.4

610

7.1

660

Potatoes

4.6

406

5.8

457

6.9

553

8.1

584

6.9b

533b

Truck vegetables

4.1

305

4.6

356

5.1

406

5.6

457

6.3b

508b

508

b

559b

6.4

813 356b

Melons

4.1

381

4.6

406

5.1

457

5.6

Strawberries

4.6

457

5.1

508

5.6

559

6.1

610

6.6

660

Citrus

4.1

508

4.6

559

5.1

660

—

—

5.6

711

Citrus (w=cover)

5.1

635

5.6

711

6.4

813

—

—

6.9

889

Dec orchard

3.8

483

4.8

533

5.8

584

6.6

635

7.6

762

Dec orchard (w/cover)

5.1

635

6.4

711

7.6

813

8.9

914

10.2

1016

Vineyards

3.6

356

4.1

406

4.8

457

5.6

508

6.4

610

a

Winter planting. Fall or winter planting. (From Ref.[2].) b

amount of water can be stored prior to the critical peak-use periods. By drawing on this stored water, peak system delivery requirements can be reduced without reducing yield potential providing the quality of the irrigation water is high.

irrigation water. By applying more water than the plants consume, most of the salts can be pushed or leached below the root zone. The first step in computing the additional water required for leaching is to determine the leaching requirement by:

Salinity Control

LR ¼

ECw 5ECe ECw

ð3Þ

All irrigation water contains some dissolved salts that are pushed downward by rainfall and application of Table 4 Guide for selecting management-allowed deficit (MAD), values for various crops MAD (%)

Crop and root depth

25–40

Shallow-rooted, high-value fruit, and vegetable crops

40–50

Orchards,a vineyards, berries, and medium-rooted row crops

50

Forage crops, grain crops, and deep-rooted row crops

a

Some fresh fruit orchards require lower multiwavelength anomalous diffraction or values during fruit finishing for sizing. (From Ref.[2].)


where LR ¼ Leaching requirement ratio for sprinkle or surface irrigation. ECw ¼ Electrical conductivity of the irrigation water [dS=m (mmhos=cm)]. ECe ¼ Estimated electrical conductivity of the average saturation extract of the soil root zone profile for an appropriate yield reduction [dS=m (mmhos=cm)]. Unless more specific information on specific crop cultivars is available the ECe values presented in Table 5, which are taken from,[3] can be used in Eq. (3). These are values that will give an approximate 10% yield reduction, as presented by.[3] (For conversion purposes: 1.0 ppm ¼ 640 EC in dS/cm.)

930

Irrigation

Table 5 Values of ECe that will give 10% yield reduction for various cropsa Crop Field crops Barley Cotton Sugar beets Wheat Soybean Sorghum Groundnut Rice Corn Flax Broadbeans Cowpeas Beans

ECe dS/m 10 9.6 8.7 7.4 5.5 5.1 3.5 3.8 2.5 2.5 2.6 2.2 1.5

Fruit and nut crops Date palm Fig, olive Pomegranate Grapefruit Orange Lemon Apple, pear Walnut Peach Apricot Grape Almond Plum Blackberry Boysenberry Avocado Raspberry Strawberry

6.8 3.8 3.8 2.4 2.3 2.3 2.3 2.3 2.2 2 2.5 2 2.1 2 2 1.8 1.4 1.3

Vegetable crops Beets Broccoli Tomato Cucumber Cantaloupe Spinach Cabbage Potato Sweet corn Sweet potato Pepper Lettuce Radish Onion Carrot Beans

5.1 3.9 3.5 3.3 3.6 3.3 2.8 2.5 2.5 2.4 2.2 2.1 2 1.8 1.7 1.5

Forage crops Tall wheat grass Bermuda grass Barley (hay) Rye grass

9.9 8.5 7.4 6.9 (Continued)


Table 5 Values of ECe that will give 10% yield reduction for various cropsa (Continued) Crop Crested wheat grass Tall fescue Sudan grass Wild rye grass Vetch Alfalfa Corn (forage) Berseem clover Orchard grass Clover a

ECe dS/m 6 5.8 5.1 4.4 3.9 3.4 3.2 3.2 3.1 2.3


Under full irrigation, where LR < 0.1, the annual deep percolation losses—even in most of the least watered areas—will normally be sufficient to provide the necessary leaching. However, under deficit irrigation or when LR < 0.1, water in addition to the consumptive use should be applied or available at some time during the year to satisfy leaching requirements. The ratio of the total depth of irrigation water required with and without leaching is equal to 1=(1 LR).

Drainage Drainage of the excess water from the soil profile and conveying it to a sump for reuse or disposal is as important as irrigation. In fact natural or man-made subsurface drainage is essential for sustaining irrigated agricultural production. Without it leaching will not be possible, causing salts to accumulate until they become toxic to plant growth, and the water table will rise and literally drown the plants. Furthermore, the productive capability of the land itself may be severely damaged and require major reclamation to become productive again. Details regarding drainage needs are presented elsewhere in this encyclopedia.

Application Depth and Frequency For periodic-move, and low-frequency, continuousmove systems, such as traveling sprinklers, it is desirable to irrigate as infrequently as practical to reduce labor costs and Eq. (1) will apply. For drip solid-set and center-pivot sprinkle systems labor costs are not a major consideration and the irrigation frequency can be selected to provide the optimal environment for plant growth, water conservation, and economic production within the physical limitations of the system.

Irrigation

Systems are usually designed so that their discharge, depths of application, and irrigation frequency meet crop water requirements during the peak consumptive use period. Therefore, systems must be managed to avoid wasting water, labor, and energy, and leaching nutrients from the soil during periods of the crops’ growth cycle when water requirements are less, when the crops’ roots may not have penetrated to their full depth, and during rainy periods. For optimum efficiency, irrigation applications should be scientifically scheduled from water budgets based on crop evapotranspiration estimates or soil moisture observations.


931

REFERENCES 1. Soil Conservation Services (SCS), Irrigation Water Requirements. In Technical Release 21. Chapter 1, Section 15, Soil Conservation Service’s (SCS) National Engineering Handbook; U.S. Department of Agriculture: Washington, DC, 1970. 2. Keller, J.; Bliesner, R.D. Soil–Water–Plant Relations. In Sprinkle and Trickle Irrigation, 1st Ed.; Van Norstrand Reinhold: New York, 1990. 3. Ayers, R.S.; Westcott, D.W. Salinity Problems. In Water Quality for Agricultures; Irrigation and Drainage Paper 29; Rev. 1, Food and Agricultural Organization of the United Nations: Rome, 1985; 31–33.

Irrigation and Soil Salinity James D. Rhoades Agricultural Salinity Consulting, Riverside, California, U.S.A.

INTRODUCTION Irrigation is an ancient practice that predates recorded history. While irrigated farmland comprises only about 15% of the worlds’ total farmland, it contributes about 36% of the total supply of food and fiber, and it stabilizes production against the vagaries of weather.[1] In 30 yr, time, irrigated agriculture is expected to have to supply 50% of the worlds’ food production requirements.[1] However, over the last 20 yr, irrigation growth has actually slowed to a rate that is now inadequate to keep up with the projected expanding food requirements.[1] Furthermore, irrigation has resulted in considerable salination of associated land and water. It has been estimated variably that the salinized area is as low as 20 and as high as 50% of the worlds’ irrigated land.[2–4] Worldwide, about 76.6 Mha of land have become degraded by human-induced salination over the last 45–50 yr.[3] It has been estimated that the world is losing at least three hectares of arable land every minute to soil salination (about 1.6 Mha per year), second only to erosion as the leading worldwide cause of soil degradation.[5–7] These data imply that the rate of salinization in developed irrigation projects now exceeds the rate of irrigation expansion.[8] Surviving the salinity threat requires that the seriousness of the problem be recognized more widely, the processes contributing to salination of irrigated lands be understood, effective control measures be developed and implemented that will sustain the viability of irrigated agriculture, and that practical reclamation measures be implemented to rejuvenate the presently degraded lands.[9,10]

DELETERIOUS EFFECTS OF SALTS ON PLANTS, SOILS, AND WATERS Salt-affected soils have reduced value for agriculture because of their content and proportions of salts, consisting mainly of sodium, magnesium, calcium, chloride, and sulfate and secondarily of potassium, bicarbonate, carbonate, nitrate, and boron. Saline soils contain excessive amounts of soluble salts for the practical and normal production of most agricultural crops. Sodic soils are those that contain excessive 932 Copyright © 2006 by Taylor & Francis

amounts of adsorbed sodium in proportion to calcium and magnesium, given the salinity level of the soil water. An example of a salt-affected irrigated soil is shown in Fig. 1 Soluble salts exert both general and specific effects on plants, both of which reduce crop yield.[11] Excess salinity in the seedbed hinders seedling establishment and in the crop root zone causes a general reduction in growth rate. In addition, certain salt constituents are specifically toxic to some plants. For example, boron is highly toxic to susceptible crops when present in the soil water at concentrations of only a few parts per million. In some woody crops sodium and chloride may accumulate in the tissue over time to toxic levels. These toxicity problems are, however, much less prevalent than is the general salinity problem. Salts may also change soil properties that affect the suitability of the soil as a medium for plant growth.[12] The suitability of soils for cropping depends appreciably on the readiness with which they conduct water and air (permeability) and on their aggregate properties (structure), which control the friability (ease with which crumbled) of the seedbed (tilth). In contrast to saline soils, which are well aggregated and whose tillage properties and permeability to water and air are equal to or higher than those of similar nonsaline soils, sodic soils have reduced permeabilities and poor tilth. These problems are caused by the swelling and dispersion of clay minerals and by the breakdown of soil structure (slaking and crusting), which results in loss of permeability and tilth. Sodic soils are generally less extensive but more difficult to reclaim than saline soils. Beneficial use of water in irrigation consists of transpiration and leaching for salinity control (the leaching requirement). Plant growth is directly proportional to water consumption through transpiration.[13] From the point of view of irrigated agriculture, the ultimate objective of irrigation is to increase the amount of water available to support transpiration. Salts reduce the fraction of water in a supply (or in the soil profile) that can be consumed beneficially in plant transpiration.[14] In considering the use of a saline water for irrigation and in selecting appropriate policies and practices of irrigation and drainage management, it is important to recognize that the total volume of a saline water supply cannot be consumed beneficially in crop production (i.e., transpired by the plant). A plant will Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001747 Copyright # 2006 by Taylor & Francis. All rights reserved.

Irrigation and Soil Salinity

Fig. 1 Photograph of salt-affected irrigated field.

not grow properly when the salt concentration in the soil water exceeds some limit specific to it under the given conditions of climate and management.[11] This is even true for halophytes.[15] Thus, the practice of blending or diluting excessively saline waters with good quality water supplies should be undertaken only after consideration is given to how it affects the volumes of consumable (usable) water in the combined and separated supplies.[14]

CAUSES OF SALINATION INDUCED BY IRRIGATION AND DRAINAGE While salt-affected soils occur extensively under natural conditions, the salt problems of greatest importance to agriculture arise when previously productive soils become salinized as a result of agricultural activities (the so-called secondary salination). The extent and salt balance of salt-affected areas has been modified considerably by the redistribution of water (hence salt) through irrigation and drainage. The development of large-scale irrigation and drainage projects, which involves diversion of rivers, construction of large reservoirs, and irrigation of large landscapes, causes large changes in the natural water and salt balances of entire geohydrologic systems. The impact of such developments can extend well beyond that of the immediate irrigated area. Excessive water diversions and applications are major causes of soil and water salination in irrigated lands. It is not unusual to find that less than 60% of the water diverted for irrigation is used in crop transpiration.[9] This implies that about 40% of the irrigation water eventually ends up as deep percolation. This drainage water contains more salt than that added with the irrigation water because of salt dissolution and mineral weathering[14] within the root zone. It often gains additional salt-load as it dissolves salts


933

of geologic origin from the underlying substrata through which it flows in its down-gradient path. This drainage water often flows laterally to lower lying areas, eventually resulting in shallow saline groundwaters of large areas of land (waterlogging). Salination occurs in soils underlain by saline shallow groundwater through the process of ‘‘capillary rise’’ as groundwater (hence, salt) is driven upwards by the force of evaporation of water from the soil surface. Correspondingly, saline soils and waterlogging are closely associated problems. Seepage from unlined or inadequately lined delivery canals occurs in many irrigation projects and is often substantial. Law, Skogerboe, and Denit[16] estimated that 20% of the total water diverted for irrigation in the United States is lost by seepage from conveyance and irrigation canals. Biswas[17] estimated that 57% of the total water diverted for irrigation in the world is lost from conveyance and distribution canals. Analogous to on-farm deep percolation resulting from irrigation, these seepage waters typically percolate through the underlying strata (often dissolving additional salts in the process), flow to lower elevation lands or waters, and add to the problems of waterlogging and saltloading associated with on-farm irrigation there. A classic example of the rise in the water table following the development of irrigation has been documented in Pakistan and is described by Jensen, Rangeley, and Dieleman[9] and Ghassemi, Jakeman, and Nix.[2] The depth to the water table in the irrigated landscape located between three major river-tributaries rose from 20 to 30 m over a period of 80–100 yr, i.e., from preirrigated time (about 1860) to the early 1960s, until it was nearly at the soil-surface. In one region, the water table rose nearly linearly from 1929 to 1950, demonstrating that deep percolation and seepage resulting from irrigation were the primary causes. Ahmad[18] concluded that about 50% of the water diverted into irrigation canals in Pakistan eventually goes to the groundwater by seepage and deep percolation. The role of irrigated agriculture in salinizing soil systems has been well recognized for hundreds of years. It is of relatively more recent recognition that salination of water resources from agricultural activities is a major and widespread phenomenon of likely equal concern to that of soil salination. The causes of water salination are essentially the same as those of soils, only the final reservoir of the discharged salt-load is a water supply in the former case.[14] The volume of the water supply is reduced through irrigation diversions and irrigation; thus, its capacity to assimilate such received salts before reaching use-limiting levels is reduced proportionately. Only in the past 15 yr has it become apparent that trace toxic constituents, such as selenium, in agricultural drainage waters can also cause serious pollution problems.[19]

934

IRRIGATION AND DRAINAGE MANAGEMENT TO CONTROL SOIL SALINITY The key to overall salinity control is strict control that maintains a net downward movement of soil water in the root zone of irrigated fields over time while minimizing excess irrigation diversions, applications, and deep percolation.[20] The direct effect of salinity on plant growth is minimized by maintaining the soilwater content in the root zone within a narrow range at a relatively high level, while at the same time avoiding surface-ponding and oxygen deletion and minimizing deep percolation. Combined methods of pressurized, high-frequency irrigation and irrigation scheduling have been developed that permit substantially the desired control to be achieved.[21,22] These systems transfer control of water distribution and infiltration from the soil to the irrigation equipment. This results in less excess water (and hence, less salt) being applied overall to the field to meet the needs of a part of the field area having lowest intake rate, as done in the more traditional gravity irrigated systems. However, gravity irrigation systems can be designed to achieve good irrigation efficiency and salinity control, even though surface ponding still does occur. The so-called level-basin, multi-set, cablegation, surge, and tailwater-return systems are among them.[21,22] The need for irrigation and the amount required to meet evapotranspiration and leaching requirement is determined from plant stress measurements, calculations of evapotranspiration amounts, measurements of soil-water depletion, measurements of soil (or soilwater) salinity, or a combination of them.[21,22] In addition to effective methods of irrigation scheduling and application, appropriate irrigation and salinity management also require an effective delivery system. Delivery systems have generally been designed to provide water on a regular schedule. Efficient irrigation systems require more flexible deliveries that can provide water on demand as each crop and particular field have need of it. Delivery systems can be improved by lining the canals, by containing the water within closed conduits, and by implementing techniques that increase the flexibility of delivery. As briefly discussed earlier, irrigated agriculture is amajor contributor to the salinity of many rivers and groundwaters, as well as soils. Reducing deep percolation generally lessens the salt load that is returned to rivers or groundwater and their pollution.[14] Additionally, saline drainage waters should be intercepted before being allowed to mix with water of better quality. The intercepted saline drainage water should be desalted and reused, disposed of by pond evaporation or by injection into some suitably isolated deep aquifer, or better yet it should be used for irrigation in a situation where brackish water is appropriate. Various



irrigation and drainage strategies have been developed for minimizing the pollution of waters from irrigation and for using brackish waters for irrigation.[14,23] Desalination of agricultural drainage waters is not now economically feasible, but improved techniques for doing this exist and some are being implemented. However, more needs to be done in this regard. Traditionally, the concepts of leaching requirement and salt-balance index have been used to plan and judge the appropriateness of irrigation and drainage systems, operations and practices with respect to salinity control, water use efficiency, and irrigation sustainability. However, these approaches are inadequate. The recommended method is to monitor directly the root-zone salinity levels and distributions across fields as a means to evaluate the effectiveness of salinity, irrigation, and drainage management practices, to detect problems (current and developing), to help determine the underlying causes of problems, and to determine source areas of major water and salt-load contributions to the underlying groundwater. Theory, equipment, and practical technology have been developed for these purposes.[24] More information about irrigation and drainage management to control soil and water salinity is found elsewhere.[25–27]

REFERENCES 1. FAO. World Agriculture Toward 2000: An FAO Study; Alexandratos, N., Ed.; Bellhaven Press: London, 1988; 338 pp. 2. Ghassemi, F.; Jakeman, A.J.; Nix, H.A. Salination of Land and Water Resources. Human Causes, Extent, Management and Case Studies; CAB International: Wallingford, UK, 1995; 526 pp. 3. Oldeman, L.R.; van Engelen, V.N.P.; Pulles, J.H.M. The Extent of human-induced soil degradation. In World Map of the Status of Human-Induced Soil Degradation: An Explanatory Note; Oldeman, L.R., Hakkeling, R.T.A., Sombroek, W.G., Eds.; International Soil Reference and Information Center (ISRIC): Wageningen, The Netherlands, 1991; 27–33. 4. Adams, W.M.; Hughes, F.M.R. Irrigation development in desert environments. In Techniques for Desert Reclamation; Goudie, A.S., Ed.; Wiley: New York, 1990; 135–160. 5. Buringh, P. Food production potential of the world. In The World Food Problem: Consensus and Conflict; Sinha, R., Ed.; Pergamon Press: Oxford, 1977; 477–485. 6. Dregne, H.; Kassas, M.; Razanov, B. A new assessment of the world status of desertification. Desertification Control Bull. 1991, 20, 6–18. 7. Umali, D.L. Irrigation-induced salinity. In A Growing Problem for Development and Environment; Technical Paper; World Bank: Washington, DC, 1993. 8. Seckler, D. The New Era of Water Resources Management: From ‘‘Dry’’ to ‘‘Wet’’ Water Savings; Consultative


9.

10.

11.

12.

13.

14.

15.

16.

17.

18.

19.

Group on International Agricultural Research: Washington, DC, 1996. Jensen, M.E.; Rangeley, W.R.; Dieleman, P.J. Irrigation trends in world agriculture. In Irrigation of Agricultural Crops; American Society of Agronomy Monograph No. 30; ASA: Madison, WI, 1990; 31–67. UNEP. Saving Our Planet: Challenges and Hopes; Nairobi, United Nations Environment Program: Nairobi, Kenya, 1992; 20 pp. Maas, E.V. Crop salt tolerance. ASCE Manuals and Reports on Engineering No. 71. In Agricultural Salinity Assessment and Management Manual; Tanji, K.K., Ed.; ASCE: New York, 1990; 262–304. Rhoades, J.D. Principal effects of salts on soils and plants. In Water, Soil & Crop Management Relating to the Use of Saline Water; Kandiah, A., Ed.; FAO (AGL) Misc. Series Publication 16=90; Food and Agriculture Organization of the United Nations: Rome, 1990; 1933. Sinclair, T.R. Limits to crop yield? In Physiology and Determination of Crop Yield; Boone, K.J., Ed.; American Society of Agronomy: Madison, WI, 1994; 509–532. Rhoades, J.D.; Kandiah, A.; Mashali, A.M. The Use of Saline Waters for Crop Production; FAO Irrigation and Drainage Paper 48; FAO: Rome, Italy, 1992; 133 pp. Miyamoto, S.; Glenn, E.P.; Oslen, M.W. Growth, water use and salt uptake of four halophytes irrigated with highly saline water. J. Arid Environ. 1996, 32, 141–159. Law, J.P.; Skogerboe, G.V.; Denit, J.D. The need for implementing irrigation return flow control. p. 1–17. In Managing Irrigated Agriculture to Improve Water Quality; Proc. Math. Conf. Manag. Irrig. Agric. Improve Water Avail., Denver, CO, May 1972; Graphics Manage. Corp.: Washington, DC, 16–18. Biswas, A.K. Conservation and management of water resources. In Techniques for Desert Reclamation; Goudie, A.S., Ed.; Wiley: New York, 1990; 251–265. Ahmad, N. Planning for Future Water Resources of Pakistan. Proceedings of Darves Bornoz Spec. Conference, National Committee of Pakistan; ICID: New Delhi, India, 1986; 279–294. Letey, J.; Roberts, C.; Penberth, M.; Vasek, C. An Agriculturl Dilemma: Drainage Water and Toxics Disposal


935

20.

21.

22.

23.

24.

25.

26.

27.

in the San Joaquin Valley; Special Publication 3319; University of California: Oakland, 1986. Rhoades, J.D. Soil salinity—causes and controls. In Techniques for Desert Reclamation; Goude, A.S., Ed.; Wiley: New York, 1990; 109–134. Hoffman, G.J.; Rhoades, J.D.; Letey, J.; Sheng, F. Salinity management. In Management of Farm Irrigation Systems; Hoffman, G.J., Howell, T.A., Solomon, K.H., Eds.; ASCE: St. Joseph, MI, 1990; 667–715. Kruse, E.G.; Willardson, L.; Ayars, J. On-farm irrigation and drainage practices. In Agricultural Salinity Assessment and Management Manual; Tanji, K.K., Ed.; ASCE Manuals and Reports on Engineering No. 71; ASCE: New York, 1990; 349–371. Rhoades, J.D. Use of saline drainage water for irrigation. In Agricultural Drainage; ASA Drainage Monograph 38; Skaggs, R.W., van Schilfgaarde, J., Eds.; ASA Drainage Monograph 38; American Society of Agronomy: Madison, WI, 1999; 619–657. Rhoades, J.D.; Chanduvi, F.; Lesch, S. Soil Salinity Assessment: Methods and Interpretation of Electrical Conductivity; FAO Irrigation and Drainage Paper 57; FAO, United Nations: Rome, Italy, 1999; 152 pp. Rhoades, J.D. Use of saline and brackish waters for irrigation: implications and role in increasing food production, conserving water, sustaining irrigation and controlling soil and water degradation. Proceedings of the International Workshop on ‘‘The Use of Saline and Brackish Waters for Irrigation: Implications for the Management of Irrigation, Drainage and Crops’’ at the 10th Afro-Asian Conference of the International Committee on Irrigation and Drainage, Bali, Indonesia, July 23–24; Ragab, R., Pearce, G., Eds.; International Committee on Irrigation and Drainage: Bali, Indonesia, 1998; 261–304. Rhoades, J.D.; Loveday, J. Salinity in irrigated agriculture. In Irrigation of Agricultural Crops; Stewart, B.A., Nielsen, D.R., Eds.; Agron. Monograph. No. 30; American Society of Agronomy: Madison, Wisconsin, 1990; 1089–1142. Tanji, K.K. Nature and extent of agricultural salinity. In Agricultural Salinity Assessment and Management; Tanji, K.K., Ed.; ASCE Manuals and Reports on Engineering No. 71, ASCE: New York, 1990; 1–17.

Irrigation Efficiency Terry A. Howell United States Department of Agriculture-Agricultural Research Service (USDA-ARS), Bushland, Texas, U.S.A.

INTRODUCTION Irrigation efficiency is a critical measure of irrigation performance in terms of the water required to irrigate a field, farm, basin, irrigation district, or an entire watershed. The value of irrigation efficiency and its definition are important to the societal views of irrigated agriculture and its benefit in supplying the high quality, abundant food supply required to meet our growing world’s population. ‘‘Irrigation efficiency’’ is a basic engineering term used in irrigation science to characterize irrigation performance, evaluate irrigation water use, and to promote better or improved use of water resources, particularly those used in agriculture and turf/landscape management.[1–4] Irrigation efficiency is defined in terms of: 1) the irrigation system performance; 2) the uniformity of the water application; and 3) the response of the crop to irrigation. Each of these irrigation efficiency measures is interrelated and will vary with scale and time. Fig. 1 illustrates several of the water-transport components involved in defining various irrigation performance measures. The spatial scale can vary from a single irrigation application device (a siphon tube, a gated pipe gate, a sprinkler, a microirrigation emitter) to an irrigation set (basin plot, a furrow set, a single sprinkler lateral, or a microirrigation lateral) to broader land scales (field, farm, an irrigation canal lateral, a whole irrigation district, a basin or watershed, a river system, or an aquifer). The timescale can vary from a single application (or irrigation set), a part of the crop season (preplanting, emergence to bloom or pollination, or reproduction to maturity), the irrigation season, to a crop season, or a year, partial year (pre-monsoon season, summer, etc.), or a water year (typically from the beginning of spring snow melt through the end of

The use of trade, firm, or corporation names in this publication is for the information and convenience of the reader. Such use does not constitute an official endorsement or approval by the United States Department of Agriculture or the Agricultural Research Service of any product or service to the exclusion of others that may be suitable. This work was done by a U.S. Government employee and cannot be copyrighted. It may be reproduced freely for any noncommercial use. 936 Copyright © 2006 by Taylor & Francis

irrigation diversion, or a rainy or monsoon season), or a period of years (a drought or a ‘‘wet’’ cycle). Irrigation efficiency affects the economics of irrigation, the amount of water needed to irrigate a specific land area, the spatial uniformity of the crop and its yield, the amount of water that might percolate beneath the crop root zone, the amount of water that can return to surface sources for downstream uses or to groundwater aquifers that might supply other water uses, and the amount of water lost to unrecoverable sources (salt sink, saline aquifer, ocean, or unsaturated vadose zone). The volumes of the water for the various irrigation components are typically given in units of depth (volume per unit area) or simply the volume for the area being evaluated. Irrigation water application volume is difficult to measure, so it is usually computed as the product of water flow rate and time. This places emphasis on accurately measuring the flow rate. It remains difficult to accurately measure water percolation volumes, groundwater flow volumes, and water uptake from shallow ground water.

IRRIGATION SYSTEM PERFORMANCE EFFICIENCY Irrigation water can be diverted from a storage reservoir and transported to the field or farm through a system of canals or pipelines; it can be pumped from a reservoir on the farm and transported through a system of farm canals or pipelines; or it might be pumped from a single well or a series of wells through farm canals or pipelines. Irrigation districts often include small to moderate size reservoirs to regulate flow and to provide short-term storage to manage the diverted water with the on-farm demand. Some on-farm systems include reservoirs for storage or regulation of flows from multiple wells.

Water Conveyance Efficiency The conveyance efficiency is typically defined as the ratio between the water that reaches a farm or field Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001826 Copyright # 2006 by Taylor & Francis. All rights reserved.

Irrigation Efficiency

937

application losses to evaporation or seepage from surface water channels or furrows, any leaks from sprinkler or drip pipelines, percolation beneath the root zone, drift from sprinklers, evaporation of droplets in the air, or runoff from the field. Application efficiency is defined as Ea ¼ 100

Fig. 1 Illustration of the various water transport components needed to characterize irrigation efficiency.

and that diverted from the irrigation water source.[1,3,4] It is defined as Ec ¼ 100

Vf Vt

ð1Þ

where Ec is the conveyance efficiency (%), Vf is the volume of water that reaches the farm or field (m3), and Vt is the volume of water diverted (m3) from the source. Ec also applies to segments of canals or pipelines, where the water losses include canal seepage or leaks in pipelines. The global Ec can be computed as the product of the individual component efficiencies, Eci, where i represents the segment number. Conveyance losses include any canal spills (operational or accidental) and reservoir seepage and evaporation that might result from management as well as losses resulting from the physical configuration or condition of the irrigation system. Typically, conveyance losses are much lower for closed conduits or pipelines[4] compared with unlined or lined canals. Even the conveyance efficiency of lined canals may decline over time due to material deterioration or poor maintenance.

Vs Vf

ð2Þ

where Ea is the application efficiency, (%), Vs is the irrigation needed by the crop (m3), and Vf is the water delivered to the field or farm (m3). The root zone may not need to be fully refilled, particularly if some root zone water-holding capacity is needed to store possible or likely rainfall. Often, Vs is characterized as the volume of water stored in the root zone from the irrigation application. Some irrigations may be applied for reasons other than meeting the crop water requirement (germination, frost control, crop cooling, chemigation, fertigation, or weed germination). The crop need is often based on the ‘‘beneficial water needs.’’[5] In some surface irrigation systems, the runoff water that is necessary to achieve good uniformity across the field can be recovered in a ‘‘tailwater pit’’ and recirculated with the current irrigation or used for later irrigations, and Vf should be adjusted to account for the ‘‘net’’ recovered tailwater. Efficiency values are typically site specific. Table 1 provides a range of typical farm and field irrigation application efficiencies[6–8] and potential or attainable efficiencies for different irrigation methods that assumes irrigations are applied to meet the crop need. Storage Efficiency Since the crop root zone may not need to be refilled with each irrigation, the storage efficiency has been defined.[4] The storage efficiency is given as Es ¼ 100

Vs Vrz

ð3Þ

where Es is the storage efficiency (%) and Vrz is the root zone storage capacity (m3). The root zone depth and the water-holding capacity of the root zone determine Vrz. The storage efficiency has little utility for sprinkler or microirrigation because these irrigation methods seldom refill the root zone, while it is more often applied to surface irrigation methods.[4]

Application Efficiency Seasonal Irrigation Efficiency Application efficiency relates to the actual storage of water in the root zone to meet the crop water needs in relation to the water applied to the field. It might be defined for individual irrigations or parts of irrigations (irrigation sets). Application efficiency includes any


The seasonal irrigation efficiency is defined as Ei ¼ 100

Vb Vf

ð4Þ

938


Table 1 Example of farm and field irrigation application efficiency and attainable efficiencies Field efficiency (%)

Farm efficiency (%)

Irrigation method

Attainable

Range

Average

Attainable

Range

Average

Surface Graded furrow w=tailwater reuse Level furrow Graded border Level basins

75 85 85 80 90

50–80 60–90 65–95 50–80 80–95

65 75 80 65 85

70 85 85 75 80

40–70 — — — —

65 — — — —

Sprinkler Periodic move Side roll Moving big gun

80 80 75

60–85 60–85 55–75

75 75 65

80 80 80

60–90 60–85 60–80

80 80 70

Center pivot Impact heads w=end gun Spray heads wo=end gun LEPAa wo=end gun

85 95 98

75–90 75–95 80–98

80 90 95

85 85 95

75–90 75–95 80–98

80 90 92

Lateral move Spray heads w/hose feed Spray heads w=canal feed

95 90

75–95 70–95

90 85

85 90

80–98 75–95

90 85

Microirrigation Trickle Subsurface drip Microspray

95 95 95

70–95 75–95 70–95

85 90 85

95 95 95

75–95 75–95 70–95

85 90 85

Water table control Surface ditch Subsurface drain lines

80 85

50–80 60–80

65 75

80 85

50–80 65–85

60 70

a

LEPA is low energy precision application. (From Refs.[6,7,11].)

where Ei is the seasonal irrigation efficiency (%) and Vb is the water volume beneficially used by the crop (m3). Vb is somewhat subjective,[4,5] but it basically includes the required crop evapotranspiration (ETc) plus any required leaching water (Vl) for salinity management of the crop root zone.

Leaching requirement (or the leaching fraction) The leaching requirement,[9] also called the leaching fraction, is defined as Vd ECi ¼ Lt ¼ Vf ECd

ð5Þ

where Lr is the leaching requirement, Vd is the volume of drainage water (m3), Vf is the volume of irrigation (m3) applied to the farm or field, ECi is the electrical conductivity of the irrigation water (dS m1), and ECd is the electrical conductivity of the drainage water (dS m1). The Lr is related to the irrigation application efficiency, particularly when drainage is the primary


irrigation loss component. The Lr would be required ‘‘beneficial’’ irrigation use V1 LT Vi ; so only Vd greater than the minimum required leaching should reduce irrigation efficiency. Then, the irrigation efficiency can be determined by combining Eqs. (4 and 5).

Vb þ LT Ei ¼ 100 Vf

ð6Þ

Burt et al.[5] defined the ‘‘beneficial’’ water use to include possible off-site needs to benefit society (riparian needs or wildlife or fishery needs). They also indicated that Vf should not include the change in the field or farm storage of water, principally soil water but it could include field (tailwater pits) or farm water storage (a reservoir) that was not used within the time frame that was used to define Ei.


939

IRRIGATION UNIFORMITY The fraction of water used efficiently and beneficially is important for improved irrigation practice. The uniformity of the applied water significantly affects irrigation efficiency. This uniformity is a statistical property of the applied water’s distribution. This distribution depends on many factors that are related to the method of irrigation, soil topography, soil hydraulic or infiltration characteristics, and hydraulic characteristics (pressure, flow rate, etc.) of the irrigation system. Irrigation application distributions are usually based on depths of water (volume per unit area); however, for microirrigation systems they are usually based on emitter flow volumes because the entire land area is not typically wetted.

Christiansen’s Uniformity Coefficient Christiansen[10] proposed a coefficient intended mainly for sprinkler systems based on the catch volumes given as CU ¼ 100

1 ð

P jÞ jX x P X

ð7Þ

where CU is the Christiansen’s uniformity coefficient (%), X is the depth (or volume) of water in each of the equally spaced catch containers (mm or ml), and x is the mean depth (volume) of the catch (mm or ml). For CU values >70%, Hart[11] and Keller and Bliesner[8] presented " CU ¼ 100 1

# s 2 0:5 x

p

ð8Þ

where s is the standard deviation of the catch depth (mm) or volume (ml). Eq. (8) approximates the normal distribution for the catch amounts. The CU should be weighted by the area represented by the container[12] when the sprinkler catch containers intentionally represent unequal land areas, as is the case for catch containers beneath a center pivot. Heermann and Hein[12] revised the CU formula (Eq. 8) to reflect the weighted area, particularly intended for a center pivot sprinkler, as follows:

CUðH&HÞ

P 39 8 2 P

> PVi Si

> > > Si Vi < Si 6 7= 7 P ¼ 100 1 6 4 5> > ðVi Si Þ > > ; : ð9Þ


where Si is the distance (m) from the pivot to the ith equally spaced catch container and Vi is the volume of the catch in the ith container (mm or ml). Low-Quarter Distribution Uniformity The distribution uniformity represents the spatial evenness of the applied water across a field or a farm as well as within a field or farm. The general form of the distribution uniformity can be given as

DUp

Vp ¼ 100 Vf

ð10Þ

where DUp is the distribution uniformity (%) for the lowest p fraction of the field or farm (lowest one-half p ¼ 1=2, lowest one-quarter p ¼ 1=4), V p is the mean application volume (m3), and V p is the mean application volume (m3) for the whole field or farm. When p ¼ 1=2 and CU > 70%, then the DU and CU are essentially equal.[13] The USDA-NRCS (formerly, the Soil Conservation Service) has widely used DUlq (p ¼ 1=4) for surface irrigation to access the uniformity applied to a field, i.e., by the irrigation volume (amount) received by the lowest one-quarter of the field from applications for the whole field. Typically, DUp is based on the post-irrigation measurement[5] of water volume that infiltrates the soil because it can more easily be measured and better represents the water available to the crop. However, the postirrigation infiltrated water ignores any water intercepted by the crop and evaporated and any soil water evaporation that occurs before the measurement. Any water that percolates beneath the root zone or the sampling depth will also be ignored. The DU and CU coefficients are mathematically interrelated through the statistical variation (coefficient of variation, s/ x, Cv) and the type of distribution. Warrick[13] presented relationships between DU and CU for normal, log-normal, uniform, specialized power, beta-, and gamma-distributions of applied irrigations. Emission Uniformity For microirrigation systems, both the CU and DU concepts are impractical because the entire soil surface is not wetted. Keller and Karmeli[14] developed an equation for microirrigation design as follows h i q m EU ¼ 100 1 1:27ðCvm Þn1=2 q

ð11Þ

where EU is the design emission uniformity (%), Cvm is the manufacturer’s coefficient of variability in emission

940


device flow rate (1 hr1), n is the number of emitters per plant, qm is the minimum emission device flow rate is the (1 hr1) at the minimum system pressure, and q mean emission device flow rate (1 hr1). This equation is based on the DUlq concept,[4] and includes the influence of multiple emitters per plant that each may have a flow rate from a population of random flow rates based on the emission device manufacturing variation. Nakayama, Bucks, and Clemmens[15] developed a design coefficient based more closely on the CU concept for emission device flow rates from a normal distribution given as CUd ¼ 100ð1 0:798ðCvm Þn1=2 Þ

ð12Þ

where CUd is the coefficient of design uniformity (%) and the numerical value, 0.798, is 0:5 2 p from Eq. (8). Many additional factors affect microirrigation uniformity including hydraulic factors, topographic factors, and emitter plugging or clogging.

WATER USE EFFICIENCY The previous sections discussed the engineering aspects of irrigation efficiency. Irrigation efficiency is clearly influenced by the amount of water used in relation to the irrigation water applied to the crop and the uniformity of the applied water. These efficiency factors impact irrigation costs, irrigation design, and more important, in some cases, the crop productivity. The water use efficiency has been the most widely used parameter to describe irrigation effectiveness in terms of crop yield. Viets[16] defined water use efficiency as WUE ¼

Yg ET

ð13Þ

where WUE is water use efficiency (kg m3), Yg is the economic yield (g m2), and ET is the crop water use (mm). WUE is usually expressed by the economic yield, but it has been historically expressed as well in terms of the crop dry matter yield (either total biomass or aboveground dry matter). These two WUE bases (economic yield or dry matter yield) have led to some inconsistencies in the use of the WUE concept. The transpiration ratio (transpiration per unit dry matter) is a more consistent value that depends primarily on crop species and the environmental evaporative


demand,[17] and it is simply the inverse of WUE expressed on a dry matter basis.

Irrigation Water Use Efficiency The previous discussion of WUE does not explicitly explain the crop yield response to irrigation. WUE is influenced by the crop water use (ET). Bos[3] defined a term for water use efficiency to characterize the influence of irrigation on WUE as WUE ¼

ðYgi Ygd Þ ðETi ETd Þ

ð14Þ

where WUE is irrigation water use efficiency (kg m3), Ygi is the economic yield (g m2) for irrigation level i, Ygd is the dryland yield (g m2; actually, the crop yield without irrigation), ETi is the evapotranspiration (mm) for irrigation level i, and ETd is the evapotranspiration of the dryland crops (or of the ET without irrigation). Although Eq. (14) seems easy to use, both Ygd and ETd are difficult to evaluate. If the purpose is to compare irrigation and dryland production systems, then dryland rather than nonirrigated conditions should be used. If the purpose is to compare irrigated regimes with an unirrigated regime, then appropriate values for Ygd and ETd should be used. Often, in most semiarid to arid locations, Ygd may be zero. Bos[3] defined irrigation water use efficiency as IWUE ¼

ðYgi Ygd Þ IRRi

ð15Þ

where IWUE is the irrigation efficiency (kg m3) and IRRi is the irrigation water applied (mm) for irrigation level i. In Eq. (15), Ygd may be often zero in many arid situations.

CONCLUSIONS Irrigation efficiency is an important engineering term that involves understanding soil and agronomic sciences to achieve the greatest benefit from irrigation. The enhanced understanding of irrigation efficiency can improve the beneficial use of limited and declining water resources needed to enhance crop and food production from irrigated lands.

REFERENCES 1. Israelsen, O.R.; Hansen, V.E. Irrigation Principles and Practices, 3rd Ed.; Wiley: New York, 1962; 447 pp.


2. ASCE. Describing irrigation efficiency and uniformity. J. Irrig. Drain. Div., ASCE 1978, 104 (IR1), 35–41. 3. Bos, M.G. Standards for irrigation efficiencies of ICID. J. Irrig. Drain. Div., ASCE 1979, 105 (IR1), 37–43. 4. Heermann, D.F.; Wallender, W.W.; Bos, M.G. Irrigation efficiency and uniformity. In Management of Farm Irrigation Systems; Hoffman, G.J., Howell, T.A., Solomon, K.H., Eds.; Am. Soc. Agric. Engrs: St. Joseph, MI, 1990; 125–149. 5. Burt, C.M.; Clemmens, A.J.; Strelkoff, T.S.; Solomon, K.H.; Bliesner, R.D.; Hardy, L.A.; Howell, T.A.; Eisenhauer, D.E. Irrigation performance measures: efficiency and uniformity. J. Irrig. Drain. Eng. 1997, 123 (3), 423–442. 6. Howell, T.A. Irrigation efficiencies. In Handbook of Engineering in Agriculture; Brown, R.H., Ed.; CRC Press: Boca Raton, FL, 1988; Vol. I, 173–184. 7. Merriam, J.L.; Keller, J. Farm Irrigation System Evaluation: A Guide for Management; Utah State Univ.: Logan, UT, 1978; 271 pp. 8. Keller, J.; Bliesner, R.D. Sprinkle and Trickle Irrigation; The Blackburn Press: Caldwell, NJ, 2000; 652 pp. 9. U.S. salinity laboratory staff. In Diagnosis and Improvement of Saline and Alkali Soils; Handbook 60; U.S. Govt. Printing Office: Washington, DC, 1954; 160 pp.


941

10. Christiansen, J.E. Irrigation by Sprinkling; California Agric. Exp. Bull. No. 570, Univ. of Calif.; Berkeley, CA, 1942; 94 pp. 11. Hart, W.E. Overhead irrigation by sprinkling. Agric. Eng. 1961, 42 (7), 354–355. 12. Heermann, D.F.; Hein, P.R. Performance characteristics of self-propelled center-pivot sprinkler machines. Trans. ASAE 1968, 11 (1), 11–15. 13. Warrick, A.W. Interrelationships of irrigation uniformity terms. J. Irrig. Drain. Eng., ASCE 1983, 109 (3), 317–332. 14. Keller, J.; Karmeli, D. Trickle Irrigation Design; Rainbird Sprinkler Manufacturing: Glendora, CA, 1975; 133 pp. 15. Nakayama, F.S.; Bucks, D.A.; Clemmens, A.J. Assessing trickle emitter application uniformity. Trans. ASAE 1979, 22 (4), 816–821. 16. Viets, F.G. Fertilizers and the efficient use of water. Adv. Agron. 1962, 14, 223–264. 17. Tanner, C.B.; Sinclair, T.R. Efficient use of water in crop production: research or re-search? In Limitations to Efficient Water Use in Crop Production; Taylor, H.M., Jordan, W.R., Sinclair, T.R., Eds.; Am. Soc. Agron., Crop Sci. Soc. Am., Soil Sci. Soc. Am.; Madison, WI, 1983; 1–27.

Irrigation Erosion Robert E. Sojka David L. Bjorneberg United States Department of Agriculture-Agricultural Research Service (USDA-ARS), Kimberly, Idaho, U.S.A.

INTRODUCTION Irrigation is important to global food production. About 15% of cropland[1] and 5% of food production land, which includes rangeland and permanent cropland,[2] are irrigated. However, irrigated land produces more than 30% of the world’s food,[3] which is 2.5 times as much per unit area compared with nonirrigated production.[1] In the United States, approximately 15% of the harvested cropland is irrigated, however, almost 40% of the total crop value is produced on irrigated land.[4] Although sprinkler- and drip-irrigated areas are increasing, most of the world’s irrigated land uses surface or flood irrigation. The countries with the largest irrigated areas are India—59,000,000 hectares (ha), China—52,580,000 ha, United States—21,400,000 ha, and Pakistan—18,000,000 ha.[2] These countries account for 55% of the world’s irrigated land; all other countries have less than 10 million ha each of irrigated land.[2] About 50% of the irrigated land in the United States is surface irrigated[5] although 95 to 99% of the irrigated land in India, China, and Pakistan is surface irrigated.[6] Soil erosion from irrigated fields has been discussed previously[7,8] and we focus on unique aspects of irrigation-induced soil erosion that are important when managing and simulating soil erosion on irrigated lands. Soil erosion mechanics can be divided into three components: detachment, transport, and deposition. Water droplets and flowing water detach soil particles; flowing water then transports these detached particles downstream; deposition occurs when flowing water can no longer transport the soil particles because flow rate decreases as water infiltrates or as rill slope or roughness change. Some particles are deposited within a few meters, although others are transported off the field with runoff water. These mechanisms are the same for surface irrigation, sprinkler irrigation, and rainfall; however, there are some systematic differences between irrigation and rainfall erosion, especially between surface irrigation and rainfall. SURFACE IRRIGATION Soil erosion is often a serious problem on surfaceirrigated land (Figs. 1 and 2). Erosion rates as high 942 Copyright © 2006 by Taylor & Francis

as 145 Mg=ha in 1 h[9] and 40 Mg=ha in 30 min[10] were reported in some early surface irrigation erosion studies. These extreme losses do not represent a sustained seasonal rate. Annual soil losses of 1 to 141 Mg=ha from surface irrigated fields were reported in a 1980 southern Idaho study.[11] Within-field erosion rates on the upper quarter of a furrow-irrigated field can be 10 to 30 times more than the field average erosion rate.[12] Some soil eroded from the upper end of a field is deposited on the lower end, whereas some soil leaves the field with runoff. Losing topsoil from the upper end of the field can decrease crop yields 25% compared with the lower end of the field.[13] Sediment cannot be transported without runoff. Runoff is planned with many surface irrigation schemes in order to irrigate all areas of the field adequately. Under ideal conditions, properly designed and managed sprinkler irrigation systems will not have any runoff from the irrigated area. However, economic and water supply constraints, along with variable slope and soil conditions, often force compromises in sprinkler irrigation design.

SPRINKLER IRRIGATION Runoff is rarely a problem with solid-set sprinkler irrigation systems because stationary sprinklers uniformly apply water at low rates (e.g., 2 mm=h). At the other end of the spectrum are systems with continuously moving laterals (center-pivot and lateral-move systems), which apply water to smaller areas (5–20 m wide) at higher rates than solid-set systems (e.g., 80 mm=h). Traveling lateral systems must irrigate large fields to reduce cost per unit area; this necessitates high instantaneous application rates to meet crop water requirements over the entire field. Application rates for center-pivot and lateral-move irrigation systems often exceed the soil infiltration rate, therefore runoff is almost always a potential problem. Sprinkler type, nozzle pressure, and nozzle size influence runoff and soil erosion by affecting application rate, wetted area, and droplet size. Low pressure sprinklers, which reduce energy costs, have smaller pattern widths and therefore greater application rates. Lower pressure also Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001575 Copyright # 2006 by Taylor & Francis. All rights reserved.

Irrigation Erosion

943

SURFACE IRRIGATION AND RAINFALL EROSION DIFFERENCES

Fig. 1 White area in the field caused by erosion from more than 80 years of surface irrigation.

produces larger drops with greater impact energy on the soil. Sprinkler systems, particularly center pivots, operate on variable slopes and topography. Slope direction relative to the lateral affects how runoff accumulates. If the lateral is perpendicular to the slope direction, runoff will tend to move away from the lateral where water is being applied, allowing water to infiltrate before traveling very far. However, if the slope is parallel to the lateral, runoff can accumulate down slope and begin flowing in erosive streams. Furthermore, if the lateral is traveling up slope, runoff will flow onto a previously wetted area; whereas with down slope travel, runoff can flow onto dry soil. These factors are further complicated by wheel tracks from moving sprinkler systems that create compacted channels for water flow.

Fig. 2 Eroded irrigation furrows near the inflow end of a field.


The most obvious difference between soil erosion from rain or sprinkler irrigation and from surface irrigation is the lack of water droplets impacting the soil during surface irrigation. This fundamental difference is important because droplet kinetic energy affects both erosion and infiltration.[14] When rain begins, droplets wet the soil surface and detach soil particles; as runoff begins, rills form in wet soil. Water flowing in rills is also exposed to falling raindrops, which affects detachment, transport, and deposition in the rills. For furrow irrigation, rills are mechanically formed in dry soil before irrigation begins. Water is applied to only a small portion of the soil surface. As water advances down the field, it flows over dry, loose soil on the first irrigation and dry, consolidated soil on subsequent irrigations. Irrigation water instantaneously wets the soil, rapidly displacing air adsorbed on internal soil particle surfaces.[15] The rapid replacement of air with water breaks apart soil aggregates,[7] increasing the erodibility of the soil. Preliminary results from a southern Idaho field study showed that soil erosion from initially dry furrows was greater than erosion from furrows that were pre-wet by drip irrigation. The hydraulics of rill flow from rain differ from furrow irrigation. Rill flow rate tends to increase downstream as additional rain water plus sheet and rill flow combine. During furrow irrigation, flow rate decreases with distance down the furrow as water infiltrates and increases with time as infiltration rate decreases which changes sediment detachment and transport capacities with distance and time. The duration of furrow irrigation runoff (typically 12 h or more) is generally longer than most rain runoff events. Temporal changes in infiltration, soil and water temperature, rill size and shape, and soil erodibility become more important for longer runoff events. Sediment concentration tends to decrease with time during furrow irrigation. Flow rate, however, increases with time, which should increase sediment detachment and transport. This indicates that soil erodibility decreases during furrow irrigation by phenomena such as armoring, surface sealing, or other unrecognized processes. Predicting small erosion events is important for irrigation. Seasonal irrigation-induced erosion occurs during numerous controlled and often small events rather than during one or two large erosion events. In southern Idaho for example, a cornfield may be sprinkler irrigated 15 to 20 times or furrow irrigated six to eight times during the growing season. The magnitude of a single irrigation erosion event is usually much smaller and less dramatic than erosion from a single 50-mm

944

thunderstorm occurring on freshly tilled soil without an established crop. However, the cumulative soil loss from irrigation during the growing season may be substantial. Chemical quality of rainfall varies less from location to location than surface water and groundwater quality. Irrigation water quality can also vary during the season as return flow is added to surface water sources or as groundwater and surface water sources are mixed. Water quality can significantly impact erosion from furrow and sprinkler-irrigated fields. Increasing electrical conductivity (EC) tends to decrease erosion whereas increasing sodium adsorption ratio (SAR) tends to increase erosion.[16,17] Interactions among EC, SAR, clay flocculation, soil chemistry, rainfall application rate, etc. influence the effects of water quality on infiltration and erosion.

REFERENCES 1. Kendall, H.W.; Pimentel, D. Constraints on the expansion of the global food supply. Ambio 1994, 23 (3), 198–205. 2. Food and Agriculture Organization. FAOSTAT– Agriculture Data. Food and Agriculture Organization On-line Database; apps.fao.org (accessed Sept. 2000). 3. Tribe, D. Feeding and Greening the World, the Role of Agricultural Research; CAB International: Wallingford, Oxon, United Kingdom, 1994. 4. National research council. In A New Era for Irrigation; National Academy Press: Washington, DC, 1996; 203 pp. 5. United States Department of Agriculture. 1998 Farm and Ranch Irrigation Survey. National Agricultural Statistics Service; www.nass.usda.gov=census= (accessed Sept. 2000).


Irrigation Erosion

6. Food and Agriculture Organization. AQUASTAT— Country Profiles. Food and Agriculture Organization Online Database; www.fao.org=WAICENT=FAOINFO= AGRICULT=AGL=AGLW=aquastat (accessed Sept. 2000). 7. Carter, D.L. Soil erosion on irrigated lands. In Irrigation of Agricultural Crops; Agronomy Monograph no. 30; Stewart, B.A., Nielson, D.R., Eds.; Am. Soc. Agronomy: Madison, WI, 1990; 1143–1171. 8. Koluvek, P.K.; Tanji, K.K.; Trout, T.J. Overview of soil erosion from irrigation. J. Irr. Drain. Eng. 1993, 119 (6), 929–946. 9. Israelson, O.W.; Clyde, G.D.; Lauritzen, C.W. Soil Erosion in Small Irrigation Furrows; Bull. 320 Utah Agr. Exp. Sta.: Logan, UT, 1946. 10. Mech, S.J. Effect of slope and length of run on erosion under irrigation. Agr. Eng. 1949, 30, 379–383. 11. Berg, R.D.; Carter, D.L. Furrow erosion and sediment losses on irrigated cropland. J. Soil Water Cons. 1980, 35 (6), 267–270. 12. Trout, T.J. Furrow irrigation erosion and sedimentation: on-field distribution. Trans. of the ASAE 1996, 39 (5), 1717–1723. 13. Carter, D.L.; Berg, R.D.; Sanders, B.J. The effect of furrow irrigation erosion on crop productivity. Soil Sci. Soc. Am. J. 1985, 49 (1), 207–211. 14. Thompson, A.L.; James, L.G. Water droplet impact and its effect on infiltration. Trans. of the ASAE 1985, 28 (5), 1506–1510. 15. Kemper, W.D.; Rosenau, R.; Nelson, S. Gas displacement and aggregate stability of soils. Soil Sci. Soc. Am. J. 1985, 49 (1), 25–28. 16. Lentz, R.D.; Sojka, R.E.; Carter, D.L. Furrow irrigation water-quality effects on soil loss and infiltration. Soil Sci. Soc. Am. J. 1996, 60 (1), 238–245. 17. Kim, K.-H.; Miller, W.P. Effect of rainfall electrolyte concentration and slope on infiltration and erosion. Soil Technology 1996, 9, 173–185.

Irrigation: Historical Perspective Robert E. Sojka David L. Bjorneberg J. A. Entry United States Department of Agriculture-Agricultural Research Service (USDA-ARS), NWISRL, Kimberly, Idaho, U.S.A.

INTRODUCTION Irrigation can be broadly defined as the practice of applying additional water (beyond what is available from rainfall) to soil to enable or enhance plant growth and yield, and, in some cases, the quality of foliage or harvested plant parts. The water source could be groundwater pumped to the surface, or surface water diverted from one position on the landscape to another. Development of irrigation water often entails development of large-scale, geographically significant dams and water impoundments and=or diversions that can provide additional functions apart from crop growth enhancement, e.g., flood control, recreation, or generation of electricity. In many cases sustainable irrigation development requires concomitant development of surface and=or subsurface drainage.

ANCIENT ORIGINS AND IMPORTANCE Irrigation may be the single most strategically important, intentional, environmental modification humans have learned to perform. While irrigation’s impact has not always been as critical to the global agricultural economy and food supply as it is today, it has always had major local impacts and profound historical and social consequences. In the Bible’s book of Genesis, we are told that God’s creation of humans was accompanied shortly thereafter by His assignation to Adam of the stewardship of the irrigated orchard that was Paradise. The four life-giving water heads of Judeo-Christian Paradise are also mentioned in the 47th Sura of the Koran.[1] Some anthropologists and historians point to the development of irrigation as the catalyst for the interaction of engineering, organizational, political, and related creative or entrepreneurial skills and activities which produced the outcome referred to as ‘‘civilization.’’[2–5] In the ancient Persian language, the word ‘‘abadan,’’ civilized, is derived from the root word ‘‘ab,’’ water.[1] Fundamental differences in social, cultural, religious, political, esthetic, economic, technological, and environmental outlook have been attributed to modern groupings of humankind related to their use of irrigation.[5] Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042707 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

The earliest archeological evidence of irrigation in farming dates to about 6000 B.C. in the Middle East’s Jordan Valley.[1] It is widely believed that irrigation was being practiced in Egypt at about the same time,[6] and the earliest pictorial representation of irrigation is from Egypt around 3100 B.C.[1] In the following millennia, irrigation spread throughout Persia, the Middle East and westward along the Mediterranean. In the same broad time frame, irrigation technology sprang up more or less independently across the Asian continent in India, Pakistan, China, and elsewhere. In the New World the Inca, Maya, and Aztec made wide use of irrigation. The technology migrated as far North as the current southwestern U.S.A., where the Hohokam built some 700 miles of irrigation canals in what is today called as central Arizona to feed their emerging civilization, only to mysteriously abandon it in the 14th century A.D.[3] In the ancient world, the level of irrigation sophistication varied from one setting to the next. The differences, however, stemmed mostly from variations in understanding of both large- and small-scale hydraulic principles, as well as the capabilities to construct feats of hydraulic engineering. The Assyrians, for example, built an inverted siphon into the Nineveh Aqueduct 700 years before the birth of Christ, an engineering feat unrivaled until the 1860 construction of the pressurized siphons of the New York Aqueduct.[3] Some ancient irrigation schemes have survived to the present day where geologic, soil, and climatic conditions were favorable and where then-known management principles were adequate for the prevailing conditions. However, some ancient schemes failed. In the Mesopotamian Valley, Syria, Egypt, and other areas throughout the Middle East, there were many cases where the principles of salt management and drainage were insufficiently understood, resulting in eventual permanent impairment of the land.[1] Siltation of ancient dams and reservoirs is a testament to inadequate soil conservation measures that eventually reduced the productivity of the land as well as destroyed the capacity of reservoirs to provide an adequate supply of water.[3] Erosion of irrigation channels, in geologically unstable areas like the Chilean deserts, and catastrophic failure of irrigation channels after 945

946

earthquakes often defeated the best efforts of ancient engineers to maintain water supplies.[3] Modern irrigation technology probably began with the Mormon settlement of the Utah Great Salt Lake Basin in 1847, and their eventual cultivation of nearly 2.5 million ha irrigated across the intermountain western U.S.A. by the turn of the century. Whereas relationships of mass, energy, and turbulence of flow were mastered at remarkably high levels of proficiency in ancient cultures, understanding of chemistry and physico-chemical interactions of soil and salt-bearing water was relatively meager even into the 19th century.

MODERNIZATION OF IRRIGATION The mid-19th century marked a conjunction of several ascending areas of scientific learning, including chemistry, physical chemistry, physics, mineralogy, and biology. These were adapted, blended, and applied in important emerging new subdisciplines of soil chemistry, soil physics, plant physiology, and agronomy, whose fundamental principles were to prove essential for sustainable irrigation system design and operation. In ancient irrigation developments, soils, climate, and water quality were used in more forgiving combinations at some locations than at others. Where seasonal rains provided leaching, where soils were permeable and well drained, and=or where irrigation water had favorable combinations of electrolyte concentrations and specific cations, irrigation has continued to the present day, even without sophisticated management. In other areas, salinization, increased soil sodicity, and elevated water tables have limited the life spans of irrigation schemes or impaired their productivity. As irrigation moved into more marginal settings, with less productive soils, poorer drainage, and greater salinity and sodicity problems, the success or failure and ultimate longevity of the schemes became more dependent on knowledgeable application and adaptation of scientific principles. America’s Mormon pioneers, choosing to settle in a remote saltimpaired desert habitat, were forced of necessity to use trial and error and the enlightened application of all available new knowledge to reclaim their lands from the desert and to practice a sustainable irrigated crop husbandry. They were so successful in their efforts that their approaches to irrigation and saltthreatened arid land reclamation and management provided the guiding principles for development of irrigation throughout the western U.S.A. from 1902 (with passage of the Reclamation Act) to the close of the 20th century.[3] The science of irrigated agriculture and arid zone soil science, in general, relied mostly on the foundation and contributions stemming from


Irrigation: Historical Perspective

these mid-19th century origins.[7] Development of irrigation in the western U.S.A. was further spurred by passage of the Desert Land Act of 1877 and the Carey Act of 1894, which provided land for settlement and governmental infrastructure for development. The first university level irrigation course is believed to have been taught by Elwood Mead (Lake Mead’s namesake) at the Agricultural College of Colorado in Fort Collins, Colorado.[8] Mead later took positions with the United State Department of Agriculture and eventually was a commissioner for the Bureau of Reclamation. Worldwide, many of the practical modern principles of irrigation system design and irrigated soil management can be traced to the lessons learned in the settling of the American west from 1847 to the close of World War II, when the total irrigated area in U.S.A. had grown to 7.5 million ha.[6] Following World War II, irrigation development worldwide entered a heady period of rapid expansion. World populations were increasing, in part because of increased life expectancies resulting from new medicines and use of dichlorodiphenyltrichloro ethane (DDT) to control malaria and other disease carrying insects. The advances in technology spurred by the first and second world wars were being applied to all avenues of life including agriculture. Electrical, steam, and internal combustion power sources became available to pump and pressurize water. New pump designs, the patenting of the center pivot and other sprinkler delivery systems came together in a few short decades between and immediately following the wars to revolutionize the ability to deliver water.[7]

CURRENT STATUS In the U.S.A., Soviet Union, Australia, and Africa huge government-sponsored programs were initiated in the 1930s, 1940s, and 1950s to build dams for hydropower, flood control, and irrigation, and to encourage settlement and stabilization of sparsely populated frontiers. The worldwide total irrigated area was about 94 million ha in 1950 and grew to 198 million ha by 1970.[9] In contrast, the world total irrigated area grew to only about 220 million ha by 1990[9] and to 263 million ha by 1996.[10] Not surprisingly, the easiest, least technically challenging, and least expensive irrigation developments occurred first, and more difficult, more technically challenging, more expensive projects dominate the remaining potential for water development. In some instances, dams, and large-scale water development projects have been hampered by poor economies and the instability of the countries in the potential development areas, rather than by the cost or technical challenges per se.


Today 60% of the earth’s grain production and half the value of all crops harvested result from irrigation.[10] Perhaps most remarkable is the agricultural production efficiency that irrigation provides worldwide. Some 50 million ha of the earth’s most productive irrigated cropland (4% of the earth’s total cropland) produces a third of the entire planet’s food crop.[11] Hectare for hectare, irrigated land produces two to two and a half times the yield and three times the crop value per hectare compared with nonirrigated land.[10,12,13] Yet, the irrigated portion amounts to only about one sixth of the world’s total cropped area[14] and about 5% of the world’s total production area, which includes cropland, range, and pasture.[15] In U.S.A., most fresh fruits and vegetables in grocery stores come from irrigated agriculture. Beyond survival and economic impact, even our entertainment and esthetics rely heavily on irrigation. Nearly allgarden nursery stock in U.S.A. is propagated and maintained under irrigation and today’s parks, play fields, golf courses, and commercial landscaping are seldom established and maintained without irrigation. To put the global production impact of irrigated agriculture in perspective, it would require over quarter billion hectares of new rainfed agricultural land (an area the size of Argentina) to supply the average additional production that irrigation’s high yield and efficiency provides. Actually, this estimate is conservative. If the land currently irrigated was no longer irrigated but left in production, its output would be well below the mean of existing rainfed land; this is because the lion’s share of irrigation occurs in arid or semiarid environments. Furthermore, additional rainfed land brought into production to replace irrigated agriculture would be well below the current rainfed average productivity; this is because the rainfed land with greatest yield potential has already been brought into production. A more realistic estimate might be double or triple the quarter billion hectare nominal replacement estimate. In a world of six billion people, irrigation has become essential by providing yet another benefit that cannot be immediately quantified, but which is as important as or more important than production efficiency or economic gain, or even the often uncredited benefits in many irrigation development schemes of hydropower, flood control, transportation, and rural development. The overriding benefit is security— security derived from food production stability. Substantial portions of the world food supply are subject to precipitous and often unpredictable yield reductions owing to drought. Irrigation was a key component of the ‘‘Green Revolution’’ of the 1960s and 1970s, which stabilized food production in the developing world, providing a new tier of nations the opportunity to turn some of their monetary and human resources to


947

nonagricultural avenues of economic and social development. Much of the drop in the rate of increase of worldwide food production in the last two decades relates to the decrease in the rate of irrigation development since 1980.

ISSUES AFFECTING THE FUTURE Although there are large projects currently underway or planned for the near future, notably in China, Pakistan, Brazil, Canada, Spain, and Portugal, the equal of the great dam building era from 1930 to 1970 will likely never be seen again. Much of the development of irrigation in the last decade has been achieved through exploitation of groundwater or by smaller scale entrepreneurial surface water developments. In Australia, for example, with the disastrous deflation of the world wool market in the 1990s, substantial numbers of individual sheep stations ceased raising animals and developed their surface water supplies to grow vast hectares of irrigated cotton and rice. Worldwide, further expansions in irrigated area are unlikely to be large because of the limited remaining surface water sources to exploit and because of the growing environmental concerns, especially related to soil waterlogging, salinization, and sodication problems. Future increases in irrigated area will likely result mainly from the development of the so-called ‘‘supplemental’’ irrigation in humid rainfed areas, from improvements in water use efficiencies associated with utilization of existing irrigation resources, and from improvements in the reuse of municipal, industrial, and agricultural wastewaters. Howell[10] noted that improved efficiencies have resulted in a reduction in the mean applied depth of water in the U.S.A. from about 650 mm annually in 1965 to 500 mm currently. These increased efficiencies have come in great part from the improved understanding of the energy physics of water which led to modern evapotranspiration (ET) theory and ET-based crop irrigation scheduling.[9,10] Many other water conservation practices were developed in the last half of the 20th century, including drip and microirrigation, which have spread from the hyperxeric conditions of Israel in the early 1950s[1] to nearly every climate and rainfall environment where there is a need, for one reason or another, to conserve water. Loss of productive capacity caused by soil salinization, sodication, and waterlogging, as well as by runoff contamination, riparian habitat impairment, and species losses, are often cited by critics of irrigation as evidence of fundamental drawbacks of irrigated agriculture. Surveys have indicated that of the existing irrigated lands, some 40–50 million ha show measurable degradation from waterlogging, salinization, and sodication.[16,17] Erosion and sedimentation of reservoirs

948

and channels caused the failures of ancient irrigation schemes and have limited the life expectancy of some modern dams to only a few decades as well.[3,18] These problems should not be trivialized. They demonstrate the need for intensified research and conservation, as well as improved dissemination and use of known prophylactic and remedial technologies. However, neither should they be overstated nor presented without due consideration of mitigating factors. If rates of production loss from these problems are weighted by relative yield or economic value of irrigation compared to rainfed agriculture, and if other positive effects of irrigation are considered, the relative magnitude of negative impacts of irrigation is greatly diminished. For example, runoff contamination from irrigated land would have to be three times the mean for rainfed land, on a crop value basis, or two to two and a half times the mean, on a yield basis, to be ‘‘comparable’’ to problems from nonirrigated agriculture because of the respective relative efficiencies of irrigated agriculture. Both the absolute and relative areas of impaired production, plus the degree of impairment, need to be compared on a global basis to rainfed losses, as well as the potential for remediation and production expansion under either circumstance. Positive impacts of irrigation water development include many social and economic benefits such as hydropower, flood control, transportation, recreation, and rural development. Positive environmental effects result from crops, field borders, canals, ditches, and reservoirs that provide significant expansions of habitat for a variety of wildlife compared to undeveloped arid land. As with all agriculture methods in recent years, irrigated agriculture has greatly improved its ability to provide humanity’s essential needs in closer harmony with environmental needs. This remains a key priority in modern irrigated agricultural research along with continued improvement of production potential to meet the needs of a growing population. Population growth is occurring mostly in underdeveloped nations, where there is an added expectation of improved diet and standard of living. This expectation raises the need for improved production per capita above a simple linear extrapolation based on population. Only high yield intensive production from irrigated agriculture has shown the potential to meet these projected needs. The knowledge and technology exist to design and operate irrigated agricultural systems sustainably, and without environmental damage or irreversible soil impairment.[9,16,19] The problem lies in implementing known scientific principles and technologies in a timely fashion as part and parcel of irrigation project and system design and management. This is true both on a regional or project basis and at the farm or field level. Politics and economics play pivotal roles in how wellknown science and technology are applied. In this



respect irrigated agriculture is no different than the myriad manifestations of rainfed agriculture, or any other environmentally impacting activity. Because of modern political and economic considerations, there is usually great pressure, when designing and developing a large-scale irrigation project, to allocate resources for development of as many irrigated hectares as possible at the outset. This often occurs without provision of an adequate technical or social support network to the farming community making the transition to irrigated agriculture. Many schemes fail to provide sufficient financial or technical resources to install drainage systems, to educate farmers, or to include them in policy formulation. The resources are needed to help guarantee prudent water application and salinity or drainage management compatible with the social, technical, and financial capabilities of the water users. These are not failures of irrigation. They are failures of human institutions. In this respect, human political, economic, and institutional considerations rather than technical advances or water availability may represent the real challenges for irrigation in the 21st century. These obstacles must be overcome if irrigated agriculture is to provide the production advantage required to satisfy future human needs and to meet improved dietary and living standard expectations.

REFERENCES 1. Hillel, D. Rivers of Eden: the Struggle for Water and the Quest for Peace in the Middle East; Oxford University Press: New York, 1994; 355 pp. 2. Mitchell, W. The hydraulic hypothesis. Curr. Anthropol. 1973, 14, 532–534. 3. Reisner, M. Cadillac Desert: the American West and its Disappearing Water; Penguin Books: New York, 1986; 582 pp. 4. Wittfogel, K.A. Oriental Despotism: a Comparative Study of Total Power; Yale University Press: New Haven, CT, 1956. 5. Worster, D. Rivers of Empire: Water, Aridity & the Growth of the American West; Pantheon Books: New York, 1985; 402 pp. 6. Hoffman, G.J.; Howell, T.A.; Solomon, K.H. Introduction. In Management of Farm Irrigation Systems; Hoffman, G.J., Howell, T.A., Solomon, K.H., Eds.; American Society of Agricultural Engineers: St. Joseph, MI, 1990; 5–10. 7. Morgan, R.M. Water and the Land, a History of American Irrigation; The Irrigation Association: Fairfax, VA, 1993; 208 pp. 8. Heerman, D.F. Where we have been, what we have learned and where we are going. In National Irrigation Symposium, Proceedings of the 4th Decennial Symposium; Evans, R.G., Benham, B.L., Trooien, T.P., Eds.; American Society of Agricultural Engineers: St. Joseph, MI, 2000; 40–51.


9. Jensen, M.E.; Rangeley, W.R.; Dieleman, P.J. Irrigation trends in world agriculture. In Irrigation of Agricultural Crops; Stewart, B.A., Nielsen, D.R., Eds.; Agronomy Monograph 30; American Society of Agronomy: Madison, WI, 1990; 31–67. 10. Howell, T.A. Irrigation’s role in enhancing water use efficiency. In National Irrigation Symposium, Proceedings of the 4th Decennial Symposium; Evans, R.G., Benham, B.L., Trooien, T.P., Eds.; American Society of Agricultural Engineers: St. Joseph, MI, 2000; 66–80. 11. Tribe, D. Feeding and Greening the World, the Role of Agricultural Research; CAB International: Wallingford, UK, 1994; 274 pp. 12. Bucks, D.A.; Sammis, T.W.; Dickey, G.L. Irrigation for arid areas. In Management of Farm Irrigation Systems, ASAE Monograph; Hoffman, G.J., Howell, T.A., Solomon, K.H., Eds.; American Society of Agricultural Engineers: St. Joseph, MI, 1990; 499–548. 13. Kendall, H.W.; Pimentel, D. Constraints on the expansion of the global food supply. Ambio 1994, 23 (3), 198–205.


949

14. Gleick, P.H. Water in Crisis: a Guide to the World? Fresh Water Resources; Oxford University Press: New York, 1993; 473 pp. 15. Food and Agriculture Organization. FAOSTAT— Agriculture Data, Food and Agriculture Organization On-Line Database. http:==apps.fao.org (accessed September 2000). 16. Rhoades, J.D. Sustainability of irrigation: an overview of salinity problems and control strategies. Pp1–42. In CWRA 1997, Annual Conference on Footprints of Humanity: Reflections on Fifty Years of Water Resource Developments, Lethbridge, Alta, June 3–6, 1997. 17. Ghassemi, F.; Jakeman, A.J.; Nix, H.A. Salinization of Land and Water Resources, Human Causes, Extent, Management, and Case Studies; CAB International: Wallingford, U.K., 1995. 18. Fukuda, H. Irrigation in the World; University of Tokyo Press: Tokyo, 1976. 19. Sojka, R.E. Understanding and managing irrigationinduced erosion. In Advances in Soil and Water Conservation; Pierce, F.J., Frye, W.W., Eds.; Sleeping Bear Press: Ann Arbor, MI, 1998; 21–37.

ISRIC: World Soil Information David Dent ISRIC–World Soil Information, Wageningen, The Netherlands

INTRODUCTION ISRIC–World Soil Information is an independent foundation, funded by the Netherlands Government with a mandate to increase knowledge of the land, its soils in particular, and to support the sustainable use of land resources; in short, to help people understand soils. Its aims are to

Inform and educate, through the World Soil Museum.

Maintain and disseminate data for the scientific community through the ICSU World Data Centre for Soils which is the custodian of global and regional datasets, land resources maps, and reports; for many of these, ISRIC is the sole repository.

Conduct applied research on land resources and their management and to support the development of national and international policy. The institute has a tradition of welcoming guest researchers.

The multilingual staff has expertise in taxonomy of soils, soil survey, land evaluation and land use planning, soil and water conservation, soil fertility, and data management and interpretation—globally and especially in tropical regions.

HISTORY On the initiative of Professor F. A. van Baren, the International Society of Soil Science promoted the establishment of an international museum of soil standards to collect, analyze, and display the soils of the world, as depicted in the FAO-UNESCO Soil Map of the World. This was adopted by the General Council of UNESCO in 1964 and support was offered by the Government of The Netherlands. The institute was founded in 1966 with working funds provided by the Ministry of Education, Culture and Sciences of The Netherlands, and administered by the International Institute for Aerospace Survey and Earth Sciences (ITC) in Enschede. At first, facilities were provided by the University of Utrecht; in 1977, the present premises—which comprise the staff quarters, laboratory, workshop, teaching and conference facilities, and the world soils exhibition—in Wageningen 950 Copyright © 2006 by Taylor & Francis

were provided by the Netherlands Directorate General of International Cooperation. In line with its evolving mandate and activities, the institute was renamed International Soil Reference and Information Centre (ISRIC) in 1974; it became a foundation with its own statues and Board of Governors in 1995, registering the logo ISRIC–World Soil Information in 2004. The business arrangement with ITC was dissolved in 2002 and a cooperative agreement concluded with Wageningen University and Research Centre. It has a memorandum of understanding with FAO, including a joint program of work; close working relations are also maintained with UNESCO, United Nations Environment Programme (UNEP), and United Nations Convention to Combat Desertification (UNCCD).

WORLD SOIL MUSEUM The World Soil Museum is the focus of an active educational program. Groups and individuals are welcome to the well-documented, thematic exhibition of soil monoliths, representing the major soils of the world, their landscape relationships, and management. Several monoliths are on loan to museums and universities around the world (Figs. 1–4). A new initiative, the ISRIC World of Soils, makes available through the Internet a wide range of educational resources: programs, pictures, data, and text. A UNESCO Chair in Land Resources is proposed as an international focus for soils education, to include current ISRIC courses in taxonomy, soil survey, land evaluation, and land use planning, and database management and contributions from several partner institutions in The Netherlands and Switzerland. Publications program includes new and updated standard technical works—in 2005 new editions of the FAO Guidelines for soil profile description and the Booker Soil Manual—and ISRIC Reports, which are available both online and in paper copy.

WORLD DATA CENTRE FOR SOILS The World Data Centre for Soils serves the scientific community under the auspices of the International Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120041260 Copyright # 2006 by Taylor & Francis. All rights reserved.

ISRIC: World Soil Information

951

Fig. 1 Collecting a soil monolith. (View this art in color at www.dekker.com.)

Fig. 3 Teaching in the World Soil Museum. (View this art in color at www.dekker.com.)

Council for Science (ICSU). Its role is to collect, scrutinize, analyze, and disseminate data and information worldwide, in particular to ICSU programs in global change, climate, and the environment. Data from ICSU programs are maintained and made freely available; a major project is under way to digitize the datasets and maps to make them available through the Internet or on DVD. The principal holdings are

domain, compiled for global climatic change studies and backed up by a working database of more than 9500 profiles.

The Global Soil and Terrain (SOTER) datasets for South and Central America, Central and Eastern Europe, and Southern and Eastern Africa: spatial mapping units and geo-located point data at scales from 1 : 1 million to 1 : 5 million, available online.

Soil and water conservation database of the World Overview of Conservation Approaches and Technologies (WOCAT).

Monolith collection of more than 900 profiles, supported by some 5000 reference samples, fully analyzed by standard methods, representing the mapping units of the FAO-UNESCO Soil map of the world[1,2] and now being extended to encompass the groups and lower-level units of the World Reference Base for Soil Resources.[3]

ISRIC soil information system (ISIS) dataset, currently 800 profile descriptions with complete, validated, laboratory data, available online in SQL.

World Inventory of Soil Emission Potentials (WISE) dataset of 4000 profiles in the public

Fig. 2 Preparation of monolith in the workshop. (View this art in color at www.dekker.com.)


For micromorphology, the following soil collections are available:

Systematic collection of 3500 large thin sections from the ISIS profiles.

STIBOKA-Jongerius collection of 10,000 large thin sections, mainly of Dutch soils.

Schmidt–Lorenz collection of more than 15,000 small thin sections of soils, mainly from Europe, Africa, Asia, and Australia.

Fig. 4 Display of monoliths. (View this art in color at www.dekker.com.)

952


Fig. 5 Soil Map of the World. (View this art in color at www.dekker.com.)

Fig. 6 A) and B) SOTER units are tracts of land with a distinctive, often repetitive, patterns of landform, slope, parent material, and soils. (View this art in color at www.dekker.com.)



Table 1 Data held in ISIS and WISE datasets: attributes of the ISIS database No. of profiles

No. of countries

Environmental characteristics

Site characteristics

Soil morphology

Physical attributes

Chemical attributes

880

81

Physiography Geology Hydrology Climate Vegetation

Surface conditions Slope Land use Human influence Crops

Horizons Color Texture Structure Consistence Cutans Nodules Coarse fragments Biological and plant activity Porosity Pans Wageningen Permeability

Particle-size distribution Water-dispersible clay Bulk density Water retention

PH Organic C Organic N CaCO3 CaSO4 Exchangeable cations Cation exchange capacity (CEC) Exchangeable Al, H Extractable Fe, Al, Si, P soluble salts Elemental composition

Mineralogical attributes

Additional information

Clay mineralogy Sand mineralogy Heavy minerals

Thin sections Photographs Soil and soil-related maps Reports Reference soil samples Classification: local, FAO, Soil Taxonomy and World Reference Base

953


954


Fig. 7 Global SOTER coverage, December 2004. (View this art in color at www.dekker. com.)

There are excellent facilities for micromorphological examination and a common catalog is in preparation to make the collections publicly accessible. The World Data Centre maintains a systematic collection of soil maps (6000): reports, specialist texts, grey literature, and journals that hold important contributions to soil literature, especially from tropical countries (17,000 titles), and 15,000 transparencies. Valuable gifts have been received from individual researchers and universities. The library and map holdings may be searched through the Wageningen University catalog (http:==library.wur.nl=isric=).

NDVI imagery to identify and delineate hotspots for further characterization by 30 m-definition Landsat imagery and subsequent field measurements by national teams; a policy initiative, Green Water Credits—akin to the Kyoto Protocol carbon credits mechanism—to build a global facility to pay rural land managers for water management services that are at present unrecognized and unrewarded, relieving poverty and enabling food and water security for everyone; and linking the SOTER and WOCAT databases with social rules to identify best-bet management practices and provide a learning network to support their application.

APPLIED RESEARCH The institute is well known for underpinning the Soil Map of the World, the Global Assessment of Human-induced Soil Degradation (GLASOD)[4] produced for the Rio Conference in 1992, the World Reference Base for Soil Resources (since 1980), the SOTER database,[5] and for interpretations of this information for land use planning and, most recently, in terms of sources and sinks of greenhouse gases. (Figs. 5–7; Table 1). The institute is also a partner in the WOCAT (www.wocat.net). The trends in research over recent decades have been shifting away from data collection and toward putting the data to work and from solo soil science to work within multidisciplinary, international teams. An example is the recently concluded study on the Pan-European Soil Erosion Assessment.[6] Current and future priorities include the following: a quantitative Global Assessment of Land Degradation and Improvement (GLADA), using 8 km-definition satellite


REFERENCES 1. FAO-UNESCO. In Soil Map of the World, 1:5M; UNESCO: Paris, France, 1974–1981; vols. I–X. 2. FAO-UNESCO-ISRIC. In Revised Legend of the Soil Map of the World. World Soil Resources Report No. 60; FAO: Rome, 1990. 3. FAO-ISRIC-ISSS. In World Reference Base for Soil Resources. World Soil Resources Report No. 84; FAO: Rome, 1998. 4. Oldeman, L.R.; Hakkeling, R.T.A.; Sombroek, W.G. World Map of the Status of Human-induced Soil Degradation; ISRIC-UNEP: Wageningen, 1990. 5. van Engelen, V.W.P.; Wen, T.T. Global and National Soils and Terrain Digital Databases (SOTER): Procedures Manual, Revised; ISRIC: Wageningen, 1995. 6. Mantel, S.; van Lynden, G.J.; Huting, J. Pan European Erosion Assessment (PESERA), Work Package 6: Scenario analysis. Final report: April 2003–September 2003, Hungary; ISRIC–World Soil Information: Wageningen, 2003.

Jhum/Shifting Agriculture P. S. Ramakrishnan School of Environmental Sciences, Jawaharlal Nehru University, New Delhi, India

INTRODUCTION The forest farmer in the tropics has managed the traditional shifting agriculture (slash-and-burn agriculture, locally known in India as jhum), which is essentially an agroforestry system organized both in space and time for centuries. The small-scale perturbations, in the past, ensured enhanced biological diversity in the forest, with enriched crop and associated biodiversity, capitalizing on the nutrient released through slashand-burn. With increasing pressure on forest resources from outside and on population pressure from within, and the consequent declining soil fertility through land degradation, agricultural cycle has shortened. It is suggested[1] that, in the late 1980s, about 500 million people were dependent on shifting agriculture in 90 countries, covering an area of approximately 400 million hectares of tropical forest land area (Table 1). A subsequent forest resource assessment[2] found that more than 7% of the 1980 forest area underwent change during the period 1980–1990, with more than half of this change due to shifting agriculture, resulting in moderate to severe degradation. The net consequence is drastic reduction in shifting agricultural cycle, leading to: 1) drastic yield reduction; 2) reduced system stability and resilience, leading to social disruption; 3) biodiversity decline because of weed takeover, biological invasion, and=or eventual site desertification; and 4) substantial CO2 emitted into the atmosphere. All attempts made so far in finding an alternate to shifting cultivation, for want of a holistic approach in dealing with the complex issues, have had little or no impact on the farmer.[3] It is in this context that an evaluation of ecological impacts and the finding of an acceptable but sustainable solution(s) to the problem become critical.

SOIL FERTILITY AND NUTRIENT BUDGET UNDER SHIFTING AGRICULTURE Shifting agriculture is largely confined to the humid tropics of Asia, Africa, and South America, with highly variable Oxisols, Ultisols, Inceptisols, and Entisols, where the major soil constraint is the presence of some toxic chemicals with low nutrient reserves.[4] The complex tropical rain forests are often extremely Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120016599 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

fragile. First, these forests, which have developed over centuries, often grow on highly infertile soil, with biomass as the chief storage compartment for nutrients. Second, in oligotrophic areas, stability is ensured by the presence of a thick surface root mat, which picks up nutrients released from decaying leaf litter on the mineral soil and pumps these nutrients back into the living biomass before they have the chance to enter the mineral soil. Frequent and large-scale perturbations, as are occurring now, upset this delicate balance in nutrient cycling. In general, nitrogen (N) and potassium (K), being very labile, are often limiting under many situations. Phosphorus (P) deficiency is a widely reported soil constraint in tropical America, which is further compounded by high P fixation related to soil acidity. There are also other problems related to soil acidity– low availability of calcium (Ca) and magnesium (Mg), often times with aluminum (Al) and manganese (Mn) toxicities. After clear-cutting and burning of the forest, the ecosystem loses its ability to hold nutrients. Losses occur through the volatilization of carbon (C) and N during the burn. Substantial nutrient losses through wind blowing of ash, runoff, and leaching through water may all occur before adequate vegetal cover builds up, during both the cropping and fallow phases. During the cropping phase, losses also occur through the uptake and removal of nutrients through biomass that gets harvested. A reduced cycle length (5 years or less) in many parts of the world has led to a drastic decline in soil fertility under short agricultural cycles imposed on the same site over a period of time (Table 1). Largely herbaceous vegetation that develops under very short cycles of 5 years or so does not help in adequate regeneration of the lost soil fertility. The rapid regeneration of forest vegetation following clearing and burning reduces nutrient loss and allows a return to the steady state cycling characteristics of mature forests. Although farmers deal with the decline in soil fertility in different ways, a minimum of 10–15 years is required for fallow regrowth in order to recoup most of the soil fertility lost during the cropping phase (Fig. 1). The C sequestration in the system depends largely on the cycle length of these systems,[1] as also observed in the Indonesian case study.[5] The N budgeting under different cycle lengths of 15, 955

956

Jhum/Shifting Agriculture

Table 1 Forest fallows (thousands of hectares) under shifting agriculture in the late 1980s Closed forest fallows

Open forest fallows

108,600

61,650

Africa

61,700

104,350

Asia

69,250

4,000

239,550

170,000

Region South America and the Pacific

(e.g., northeastern India). The consequence of this is less land area of shifting agriculture and a further shortening of the cycle. Therefore, the success of shifting agriculture is, to a large extent, related to nutrient cycling patterns and processes of the forest fallow phase.

(From Ref. .)

KNOWLEDGE SYSTEMS FOR LAND USE MANAGEMENT

10, and 5 years is illustrative of the kind of issues involved (Table 2). During one cropping phase, the agroecosystem loses about 600 kg ha1 N (the difference between the soil N capital before and after one cropping). With the plot under a 5-year cycle having the same cycle length during the last 20 years, this system had lost 1.28 103 kg ha1 N from its initial capital of 7.68–6.40 103 kg ha1. While 10and 15-year agricultural cycles are long enough to restore the original N status in the soil before the next cropping, it seems unlikely that the 600 kg ha1 N lost during one cropping could be restored under a 5-year cycle—an observation similar to other nutrients, too, such as K. The increased frequency of fire and cropping, with too short a fallow phase, thus results in rapid site degradation. The first step in site degradation is the replacement of forests by an arrested weed stage. Large tracts of forested lands in the Asian tropics, for example, are taken over by the grass Imperata cylindrical (thatching grass, locally known in the Asian region as ‘‘alangalang’’). Exotic weed invasion is yet another major consequence of frequent perturbation under short agricultural cycles in many parts of the world, with consequences for ecosystem processes.[6] In extreme cases, the end result is a bald, totally desertified landscape[1]

‘‘Formal’’ ecological knowledge derived through the hypothetico-deductive method—validated ‘‘traditional ecological knowledge’’ (TEK) derived largely through societal experiences and perceptions accumulated by traditional societies during their interaction with nature and natural resources—has a strong human element attached to it. This knowledge needs to be effectively integrated to ensure the participatory land use development of traditional societies.[7] Linking ecological processes with social processes is the key issue here. Thus, for example, the concept of ecological keystone, an end product of a social selection process, is illustrative of the linkages that exist between the traditional and the formal knowledge systems. Thus, in many areas in northeast India where the landscape is highly degraded, a legume crop of lesser known food value, Flemingia vestita (locally known in Megalaya, India, as ‘‘Soh-phlong,’’ yielding up to 3000 kg of edible juicy tuber), which is socially valued, is used both in space and in time under a 2- to 5-year rotational fallow system. By fixing 250 kg ha 1 yr1 N, this keystone species ensures the sustainability of these low-input agroecosystems, under conditions of extreme pressure on land under low soil fertility. Using organic residues and manipulating soil biodiversity through socially valued and ecologically important keystone

Tropical world (total) [1]

Fig. 1 Changes in cumulative quantity of available P and K within a soil column of 40 cm depth under 0-, 1-, 5-, 10-, 15-, and 50-year-old jhum fallows (e.g., g eq=m2).



957

Table 2 Net change of N (103 kg ha1 yr1) in the soil under jhum in northeast India 5-year fallow cycle 15-year fallow cycle

10-year fallow cycle

First year crop

Second year crop

Soil pool before burning

7.68

7.74

6.40

5.98

Soil pool at the end of cropping

7.04

7.15

5.98

5.60

Net difference

0.64

0.59

0.42

0.38

(From Ref.[3].)

earthworm species, sustainable fertility management has been made possible. It is a patented technology that offers opportunities[8] for effective fallow management practices. Nepalese alder (Alnus nepalensis), a socially valued species, is another N-fixing tree conserved by traditional societies of northeast India in their shifting agricultural plots. It happens to be a socially significant keystone species. This early successional tree species in the northeastern hill region is conserved by the shifting agricultural farmer, during both the cropping and fallow phases. With its roots nodulated by Frankia, this species can fix up to about 125 kg ha1 yr1 N and has the potential to recover all the 600 kg of N lost from the system over a 5-year cycle period. It would, otherwise, take a minimum of 10 years of natural fallow regrowth to recover all these N back into the system. In other words, there is a close connection between the ecological and social dimensions of this validated TEK, and there are implications for fallow management through ‘‘incremental pathway’’ and=or ‘‘contour pathway’’ for agroecosystem redevelopment.[3,9] Adapting this knowledge to modern scientific inputs is important for community participation in soil fertility management and acceleration of the developmental process itself.

at least in the short term. Management of forest fallows[3] or grass fallows[10] seems to be an attractive cost-effective solution to the problem, as has been suggested through many studies. Such an approach forms the basis for a major initiative, which aims at redeveloping shifting agriculture in Nagaland in northeast India,[11] through a participatory process of fallow management in over 5500 replicated test plots in farmers’ fields spread across 1200 villages. The key to the management lies in appropriately constructed village-level institutions that are based on the local value system. In parts of South America (Venezuela), natural secondary forest succession has been suggested as a model to replace traditional shifting agriculture: bean (Phasolus vulgaris), corn (Zea mays), sugarcane (Saccharum officinarum), and pineapple (Ananas comosus) in the first year; followed by woody yucca (Manihot esculenta), cashew (Anacardium occidentale), or papaya (Carica papaya); followed by larger trees such as Brazil nut (Bertholletia excelsa) and jackfruit (Artocarpus sp.). At another level, the home garden concept could be the model for developing a plantation economy (found in many Asian and Latin American countries). Contour Pathway

SHIFTING AGRICULTURE AND SUSTAINABILITY Realizing that sustainable soil fertility management is the key issue for finding a solution to the vexed problem of shifting agriculture-affected areas, two different approaches are possible. If the society under consideration is more ‘‘traditional,’’ an incremental pathway is more appropriate to be able to relate to their value system. For others who are less traditional and are ready to make a more drastic departure, the contour pathway could be the solution. Incremental Pathway Building on traditional technology in an incremental fashion is one of the options for shifting agriculture,


As a possible medium strategy, this type of management acknowledges and works with the ecological forces that provide the base on which the system must be built, while acknowledging the social, economic, and cultural requirements of the farming communities. Working with nature instead of dominating it, this approach seeks active planning, keeping in mind the nature of the background ecosystem. Slope management has long been a major element of farming in the western Pacific region and in the uplands of the Asian tropics. The Sloping Agricultural Land Technology (SALT) developed by the Mindanao Baptist Rural Life Center in the early 1980s in the southern part of the Philippines is based on planting field and perennial crops in 3- to 5-m bands between double rows of N-fixing trees and shrubs planted on contours for soil conservation.[12] The crop species and the tree=shrub species could vary. Although SALT technology has been tested successfully

958

in many countries in the Asian tropics, socioeconomic and cultural problems stand in the way of its large-scale acceptance.

CONCLUSIONS Involving local communities in managing forestry and nontimber forest product-related activities for cash economy is important, as shifting agriculture cannot be viewed in isolation from forestry and forest-related activities of local communities. In the ultimate analysis, as a long-term strategy, it may be desirable to have a mosaic of agroecosystem types using all of the pathways just mentioned, along with intensive modern agricultural systems, coexisting with natural ecosystem types, managed or unmanaged. The maintenance of the overall sustainability of the system requires a patchwork mosaic that would, albeit inadvertently, be the best plan for effectively managing natural resources of the landscape, in the shifting agricultureaffected areas.

REFERENCES 1. FAO=UNEP. Tropical Forest Resources (by Jean-Paul Lanley); Forestry Paper 50; Food and Agriculture Organization (FAO): Rome, Italy, 1982. 2. FAO. Forest Resources Assessment 1990: Global Synthesis; Forestry Paper 124; Food and Agriculture Organization (FAO): Rome, Italy, 1995. 3. Ramakrishnan, P.S. Shifting Agriculture and Sustainable Development: An Interdisciplinary Study from North-Eastern India; UNESCO-MAB Series; Paris, Parthenon Publishers: Carnforth, Lancs, UK, 1992. 4. Sanchez, P.A. Soils. In Tropical Rain Forest Ecosystems; Leith, H., Werger, M.J.A., Eds.; Elsevier: Amsterdam, 1989; 73–88.



5. Tomich, T.P.; van Noordwijk, M.; Budidaresono, S.; Gillison, A.; Kusumanto, T.; Murdiayarso, D.; Stolle, F.; Fagi, A.M. Alternatives to Slash-and-Burn in Indonesia: Summary Report and Synthesis of Phase II; ASB Indonesia and ICRAF-S.E. Asia: Bogor, Indonesia, 1998. 6. Ramakrishnan, P.S.; Vitousak, P.M. Ecosystem-level processes and consequences of biological invasions. In Biological Invasion: A Global Perspective; Drake, J.A., Mooney, H.A., di Castri, F., Groves, R.H., Kruger, F.J., Rejmanek, M., Williamson, M., Eds.; SCOPE 37; John Wiley: New York, 1989; 281–300. 7. Ramakrishnan, P.S. Ecology and Sustainable Development; National Book Trust: New Delhi, India, 2001. 8. Senapati, B.K.; Naik, S.; Lavelle, P.; Ramakrishnan, P.S. Earthworm-based technology application for status assessment and management of traditional agroforestry systems. In Traditional Ecological Knowledge for Managing Biosphere Reserves in South and Central Asia; Ramakrishnan, P.S., Rai, R.K., Katwal, R.P.S., Mehndiratta, S., Eds.; UNESCO and Oxford IBH: New Delhi, 2002; 139–160. 9. Swift, M.J.; Vandermeer, J.; Ramakrishnan, P.S.; Anderson, J.M.; Ong, C.K.; Hawkins, B. Biodiversity and agroecosystem function. In Functional Roles of Biodiversity: A Global Perspective; Mooney, H.A., Cushman, J.H., Medina, E., Sala, O.E., Schulze, E.D., Eds.; SCOPE Series; John Wiley: Chichester, UK, 1996; 261–298. 10. Lal, R.; Wilson, G.F.; Okigbo, B.N. Changes in properties of an alfisol produced by various cover crops. Soil Sci. 1979, 127, 377–382. 11. NEPED; IRRR. Building Upon Traditional Agriculture in Nagaland; Nagaland Environmental Protection and Economic Development, Nagaland India and International Institute of Rural Reconstruction: Philippines, 1999. 12. Pratap, T.; Watson, H.R. Sloping Agricultural Land Technology (SALT): A Regenerative Option for Sustainable Mountain Farming; International Centre for Integrated Mountain Development: Khatmandu, Nepal, 1994.

Land Capability Analysis Michael J. Singer University of California, Davis, California, U.S.A.

INTRODUCTION

KINDS OF SYSTEMS

All soils are not the same and they are not of the same capability for every use. Land capability implies that the choice of land for a particular use contributes to the success or failure of that use. It further implies that the choice of land for a particular use will determine the potential impact of that use on surrounding resources such as air and water. To make the best use of land and to minimize the potential for negative impacts on surrounding lands, land capability analysis is needed. The assessment of land performance for specific purposes is land evaluation.[1] A system that organizes soil and landscape properties into a form that helps to differentiate among useful and less useful soils for a purpose is land capability classification. Land capability is a broader concept than soil quality, which has been defined as the degree of fitness of a soil for a specific use.[2] Bouma[3] points out that land capability or land potential needs to be evaluated via various scales and gives the example of precision agriculture, which requires land capability analysis in more detail than that required to determine if an investment should be made to initiate agriculture. Land capability or suitability classification systems have been designed to rate land and soil characteristics for specific uses (Table 1). Huddleston[4] has reviewed many of these systems. They may also rate land qualities. Land qualities have been defined by the United Nations Food and Agricultural Organization[1] as ‘‘attributes of land that act in a distinct manner in their influence on the function of land for a specific kind of use.’’ An example of a land quality is the plant available water stored in soil, and an example of a soil characteristic is the clay content that contributes to the plant available water holding capacity. It is generally recognized that a single soil characteristic is of limited use in evaluating differences among soils,[5] and that use of more than one quantitative variable requires a system for combining the measurements into a useful index.[6] Gersmehl and Brown[7] advocate regionally targeted systems.

Land rating systems for agriculture include those that are used to evaluate the potential for agricultural development of new areas, and others that evaluate the potential for agriculture in already developed areas. Many other land capability assessments exist to help planners rate suitability of agricultural lands for nonagricultural uses. Some examples of agricultural land capability systems are described in this chapter. All systems have, in common, a set of assumptions on which the analysis is based and each system answers the question ‘‘capability for what use.’’


Development Potential Examples of systems designed to determine the potential for agricultural development include the FAO framework for land evaluation and the U.S. Bureau of Reclamation (USBR) irrigation suitability classification. The FAO framework combines soil and land properties with a climatic resources inventory to develop an agro-ecological land suitability assessment. The USBR capability classification was frequently used to evaluate land’s potential for irrigation in the Western U.S. during the period of rapid expansion of water delivery systems.[8–11] It combines social and economic evaluations with soil and other ecological variables to determine whether the land has the productive capacity, once irrigated, to repay the investment necessary to bring water to an area. It recognizes the unique importance of irrigation to agriculture and the special qualities of soils that make them irrigable. Land Capability The USDA Land Capability Classification (see entry by Fenton) is narrower in scope than either the FAO or USBR capability rating systems. The purpose of the Land Capability Classification (LCC) is to place arable soils into groups based on their ability to sustain common cultivated crops that do not require specialized site conditioning or treatment.[12] Nonarable soils, unsuitable for long term, sustained cultivation,

959

960

Land Capability Analysis

Table 1 Examples of land capability systems System

Purpose

Property

References

FAO framework

Development potential

Used for large-scale development of agriculture

[1,10]

USBR irrigation suitability

Potential for irrigation development

Used for determining potential to repay costs of developing irrigation

[8,9]

USDA land capability classification

Land capability for agriculture

Uses 13 soil, climate, and landscape properties to determine agricultural capability

[4,12]

Storie index

Land capability for agriculture

Uses nine soil and management factors to determine agricultural capability

[17–19]

Soil potential

Soil suitability for specific uses

Uses a cost index to rate land for any potential use

[16]

Soil quality

Determining status of soil profile

Used for determining the status of selected soil properties

are grouped according to their ability to support permanent vegetation, and according to the risk of soil damage if mismanaged. Several studies have shown that lands of higher LCC have higher productivity with lower production costs than lands of lower LCC.[13–15] In a study of 744 alfalfa-, corn-, cotton-, sugar beet- and wheat-, growing fields in the San Joaquin Valley of California, those with LCC ratings between 1 and 3 had significantly lower input=output ratios than fields with ratings between 3.01 and 6.[15] The input=output ratio is a measure of the cost of producing a unit of output and is a better measure of land capability than output (yield) alone. This suggests that the LCC system provides an economically meaningful assessment of agricultural soil capability. Quantitative systems result in a numerical index, typically with the highest number being assigned to the land or soil with the highest capability for the selected use. The final index may be additive, multiplicative, or more complex functions of many land or soil attributes. Quantitative systems have two important advantages over nonquantitative systems: 1) they are easier to use with GIS and other automated data retrieval and display systems and 2) they typically provide a continuous scale of assessment.[16] No single national system is presently in use but several state or regional systems exist. One example of a quantitative system is the Storie index rating (SIR). Storie[17] determined that land productivity is dependent on 32 soil, climate, and vegetative properties. He combined only nine of these properties into the SIR, to keep the system from becoming unwieldy. The nine factors are soil morphology (A), surface texture (B), slope (C), and management factors drainage class (X1), sodicity (X2), acidity (X3), erosion (X4), micro-relief (X5), and fertility (X6). Each


[6,16,20]

factor is rated from 1 to 100%. These are converted to their decimal value and multiplied together to yield a single rating for a soil map unit[17–19] [Eq. (1)]: SIR ¼

A B C

6 Y

! Xi

100

ð1Þ

i¼1

An area-weighted SIR can be calculated by multiplying the SIR for each soil map unit within a parcel by the area of the soil unit within the parcel, followed by summing the weighted values and dividing by the total area [Eq. (2)]. 1 total area n X SIR soili area soili :

Area-weighted SIR ¼

i¼1

ð2Þ Values for each factor were derived from Storie’s experience mapping and evaluating soils in California, and in soil productivity studies in cooperation with California Agricultural Experiment Station cost-efficiency projects relating to orchard crops, grapes and cotton. Soils that were deep, which had no restricting subsoil horizons, and held water well had the greatest potential for the widest range of crops. Reganold and Singer[15] found that area-weighted average SIR values between 60 and 100 for 744 fields in the San Joaquin Valley had lower but statistically insignificant input=output ratios than fields with indices 5:0 SMB ¼ 35:3 NRN if soil pH is 5:0 Both CFI and CEF methods yield good estimates of SMB, but CHCl3 is a biohazard. Also, the CFI method is affected by high organic matter content, organic amendments, low pH, and soil waterlogged conditions, and is time-consuming and involves several steps.[1,9] The CFE method is fast and useful where CFI does not work.[6] A portion of the SMB may be insensitive to CHCl3 fumigation.

Microwave (MW) Irradiation Incubation and Extraction Methods The underlying principle of the MW irradiation method is to use nonthermal MW energy at 800 J=g oven-dried equivalent of field-moist soil in plastic tubes with punctured caps to disrupt the microbial cells and then incubate or extract the MW and unmicrowaved soils to measure flushes of C.[8] In MW irradiation incubation method, both MW and unmicrowaved soils are incubated for 10 days in the dark at 25 C, and the flush of CO2 is absorbed in dilute solution of NaOH followed by an acid–base titration to calculate the SMB. SMB ¼ CO2 CMW =KMI where CO2CMW is the net flush of CO2 from MW and unmicrowaved soils, respectively, and KMI is a coefficient of 0.341. In MW irradiation extraction method, the MW soil is extracted for flush of C by neutral 0.5 M K2SO4 compared to unmicrowaved soil.[8] An automatic analyzer with ultraviolet (UV) persulfate oxidation and infrared (IR) detection, or rapid colorimetric[8,15] or titrimetric method, is used to determine extracted C for the SMB measurement. SMB ¼ CEXTMW =KME where CEXTMW is the net flush of C from MW and unmicrowaved soils, respectively, and KME is the extraction coefficient (0.213) of the SMB.[8] The MW irradiation method is an alternate approach to CFI and CFE methods for rapid, precise, safe, and reliable measurement of SMB.[8] This method is very economical. To avoid the release of nonbiomass C from soil, the MW oven has to be calibrated before use.


Microbial Biomass Measurement Methods

Rehydration Method Rewetting of air-dry soils with dilute salt solution or water is the basis for the rehydration method for determining SMB.[16,17] Upon rehydration, the desiccated microbial cells in air-dried soils are disrupted and they release intracellular C compounds. The extracted C is analyzed by a rapid K2Cr2O7 oxidation method to determine the SMB content. SMB ¼ Ec =0:23 where Ec is the net flush of 0.25 M K2SO4 extracted C from air-dried and field-moist soils, respectively, and 0.23 is the extraction coefficient.[17] The rehydration method is a simple procedure for SMB determination that does not use hazardous chemicals, although prolonged air-drying of soil often releases nonbiomass C. A portion of the SMB may be insensitive to air-drying.

Freeze-Dried Soil Extraction Method The principle of the method is that extraction of freezedried soils with either 0.5 M K2SO4 or 0.5 M NaHCO3 releases cytoplasmic C compounds from desiccated and disrupted microbial cells.[7] The extracted C is analyzed by a rapid colorimetric method[15] to calculate the SMB contents. SMB ¼ C K2 SO4 =0:152 SMB ¼ C NaHCO3 =0:257 where C K2SO4 and C NaHCO3 are the net differences in C extracted by 0.5 M K2SO4 and NaHCO3 from freeze-dried and field-moist soils, respectively. This is a precise, reliable, and safe method for measuring SMB but it requires trained personnel and sophisticated equipment.

Adenosine Triphosphate Extraction Method The principle of this method is that, by using suitable reagents, soil microbial cells are rapidly disrupted. This is followed by stabilization of ATP with deactivating synthesis and degradative enzyme processes, and extraction of ATP from the soil matrix. Then, using luciferin–luciferase light reactive system, SMB contents can be calculated.[18,19] 1 mg ATP ¼ 250 mg of SMB As ATP is rapidly degraded during extraction and adsorbed by soil constituents, SMB measurement is


often uncertain because of storage conditions, season of collection, low and irregular recovery, and weak correlation to SMB measured.[1]

Phopholipid Fatty Acids Extraction Method This method is based on extraction of phospholipid fatty acids (PFLA) from soil microbial cell membranes via suitable extractants.[11,14] The lyophilized soil is extracted with the single-phase CHCl3–CH3OH–buffer system, followed by analysis of PFLA by a capillary gas chromatography (GC) with flame ionization detection to SMB.[14] 1 nmol PFLA/g soil ¼ 2:4 mg SMB Extracts are easily analyzed to identify the different PLFAs for the determination of SMB,[5] but this is a time-consuming, complex, and expensive method.

Substrate-Induced Respiration (SIR) Method The SIR method, which was first reported by Anderson and Domsch,[10] utilizes the initial change in the soil respiration response as a result of adding glucose or sucrose and nutrients. The SMB is calculated from the maximum initial respiratory response (MIRR) measured at 22 C, as follows: SMB ¼ ð40:04MIRRÞ þ 0:37 where MIRR is mL CO2=g soil. This is a fast method to measure SMB, but may often overestimate by measuring the glucose responsive active portion of the SMB. The use of glucose may shut down the metabolism of other microbes. The SIR method requires a GC to measure evolved CO2.

UV SPECTROSCOPIC METHOD This technique is a rapid and inexpensive method to estimate SMB based on the near-UV light absorption of certain molecules (nucleic acids=nucleotides) of microbial cells extracted from CHCl3 fumigated soils.[3] UV absorbance at 280 nm is significantly correlated to SMB measured by CHCl3 method as follows. SMB ¼ 21; 747 E280 nm Although a fast and simple method to determine SMB, this process uses CHCl3. The results are often compromised by soil colloidal interferences and electrolyte precipitation.[3]


1069

CONCLUSIONS Common and new methods were discussed, including their advantages and disadvantages. Over the years, the indirect methods have been more widely used compared to direct methods for simple, rapid, and precise measurement of SMB. Among the indirect methods, CHCl3 fumigation incubation has been used as a baseline for calibrations and correlations for other methods. Extraction methods, in particular the MW soil extraction, are gaining increasing acceptance for their simplicity, rapidity, and precision in determining SMB. No method is universally accepted because of various limitations. Consequently, there is a demand for further improvement in SMB measurement.

REFERENCES 1. Martens, R. Current methods for measuring microbial biomass-C in soil: potentials and limitations. Biol. Fertil. Soils 1995, 19, 77–81. 2. Beck, T.; Joergensen, R.G.; Kandler, E.; Makeschin, F.; Nuss, E.; Oberholzer, H.R.; Scheu, S. An interlaboratory comparison of ten different ways of measuring soil microbial biomass C. Soil Biol. Biochem. 1997, 29, 1023–1032. 3. Nunan, N.; Morgan, M.A.; Herlihy, M. Ultraviolet absorbance (280 nm) of compounds released from soil during chloroform fumigation as an estimate of the microbial biomass. Soil Biol. Biochem. 2002, 30, 1599–1603. 4. Alexander, M. Most probable number method for microbial populations. In Methods of Soil Analysis, Part II. Chemical and Microbiological Properties, 2nd Ed.; Page, A.L., Miller, R.H., Keeney, D.R., Eds.; American Society of Agronomy: Madison, WI, 1982; Vol. 9, 815–820. 5. Bottomley, P.J. Light microscopic methods for studying soil microorganisms. In Methods of Soil Analysis, Part II. Microbiological and Biochemical Properties; Weaver, R.W., Angle, J.S., Botomley, P.S., Eds.; Soil Science Society of America, Inc.: Madison, WI, 1994; 81–106. 6. Vance, E.D.; Brooks, P.C.; Jenkinson, D.S. An extraction method for measuring soil microbial biomass-C. Soil Biol. Biochem. 1987, 19, 703–707. 7. Islam, K.R.; Weil, R.R.; Mulchi, C.L.; Glenn, S.D. Freeze-dried soil extraction method for the measurement of microbial biomass C. Biol. Fertil. Soils 1997, 24, 205–210. 8. Islam, K.R.; Weil, R.R. Microwave irradiation of soil for the routine measurement of microbial biomass C. Biol. Fertil. Soils 1998, 27, 408–416. 9. Jenkinson, D.S.; Powlson, D.S. The effects of biocidal treatments on metabolism in soil. V. A method for measuring soil biomass. Soil Biol. Biochem. 1976, 8, 209–213.

1070

10. Anderson, J.P.E.; Domsch, K.H. A physiology method for quantitative measurement of microbial biomass in soils. Soil Biol. Biochem. 1978, 10, 519–525. 11. White, D.; Davis, W.M.; Nickels, J.S.; King, J.D. Determination of the sedimentary microbial biomass by extractable lipid phosphate. Oecologia 1979, 40, 51–62. 12. Webster, J.J.; Hampton, G.J.; Leach, F.R. ATP in soil: a new extractant and extraction procedure. Soil Biol. Biochem. 1984, 16, 335–342. 13. Joergensen, R.G. Quantification of the microbial biomass by determining ninhydrin-reactive N. Soil Biol. Biochem. 1996, 28, 301–306. 14. Baily, V.L.; Peacock, A.D.; Smith, J.L.; Bolton, H., Jr. Relationship between soil microbial biomass determined by chloroform fumigation–extraction, substrateinduced respiration, and phospholipid fatty acid analysis. Soil Biol. Biochem. 2002, 34, 1385–1389. 15. Islam, K.R.; Weil, R.R. A rapid microwave digestion procedure for spectrophotometric measurement of soil



16.

17.

18.

19.

organic C. Commun. Soil Sci. Plant Anal. 1998, 29, 2269–2284. Blagodatskaya, S.A.; Blagodatskaya, Y.V.; Yu, A.; Panikov, N.S. A rehydration method of determining the biomass of microorganisms in soil. Pochvovedeniye 1987, 4, 64–71. Sikora, L.J.; Yakovchenko, V.; Kaufman, D.D. Comparison of the rehydration method for biomass determination to fumigation–incubation and substrate-induced respiration method. Soil Biol. Biochem. 1994, 26, 1443–1445. Oades, J.M.; Jenkinson, D.S. Adenosine triphosphate content of the soil microbial biomass. Soil Biol. Biochem. 1979, 11, 201–204. Parkinson, D.; Paul, E.A. Microbial biomass. In Methods of Soil Analysis, Part 2: Chemical and Microbiological Properties, 2nd Ed.; Page, A.L., Ed.; American Society of Agronomy: Madison, WI, 1982; Vol. 9, 821–830.

Microbial Communities K. R. Islam S. R. Wright The Ohio State University South Centers, Piketon, Ohio, U.S.A.

INTRODUCTION The soil, a complex and heterogeneous component of terrestrial ecosystems, sustains an immense number and diversity of microorganisms, which require much more study. Currently, less than 1% of the soil microorganisms have been cultured, identified, and characterized in the soil ecosystem. Soil microorganisms, prokaryotic and eukaryotic, spend all or part of their lives in the soil environment.[1,2] These microorganisms live together by interacting with each other and with other organisms in the soil. The soil microbial community presents a complex and variable association among the different levels of biological organization, which encompasses genetic variability and the richness and relative evenness in communities. There is a growing interest in the relationships among ecosystem diversity, structure, and function. A number of theories have been formulated on how microbial species composition and diversity relate to the functional capability of the terrestrial ecosystem.[3] It has been suggested that enhanced microbial species diversity is beneficial to ecosystem function.[4,5] Others suggest that the properties of an ecosystem depend more on the functional capabilities of a particular microbial species than on the total number of species.[6–8]

COMPOSITION OF SOIL MICROBIAL COMMUNITIES The soil contains billions of organisms representing nearly every phylum.[9] Microbial communities are the most complex and diverse group of soil organisms, ranging in size from 0.5 to 5.0 mm, which consists predominantly of bacteria, fungus, actinomycetes, and lichens.[2,10] A simplified, general classification of soil microorganisms is presented in Table 1. Soil bacteria are single-celled, prokaryotic organisms in various shapes and sizes (0.5–5 mm). They are perhaps the most complex and diverse group of soil microorganisms (about 20,000 different species per gram of soil) and are adapted to most environments.[2,10] In general, bacteria are not energy-efficient Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120014249 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

organisms in soil (only about 20% of the organic substances decomposed may become bacterial tissue).[11] Soil fungi are eukaryotic organisms that may form visible macroscopic structures in soil. Fungi comprise an extremely diverse group of microorganisms with tens of thousands of species present in the soil. Although they are usually in somewhat smaller numbers than bacteria, fungi dominate the biomass and metabolic activity in many soils because of their relatively large size and branching. Fungi are energy-efficient organisms in soil (30–50% of the organic substances decomposed may become bacterial tissue).[11] Like the fungi, actinomycetes are filamentous and often profusely branched cells, although their mycelia threads are much smaller than those of fungi. Actinomycetes were previously classified with the fungi; however, they are now classified as bacteria. They have no nuclear membrane and often dissociate into spores that closely resemble bacterial cells. We will consider them separately because of the unique roles they play in the soil compared to bacteria.[2,10] Although actinomycetes develop best in moist, warm, well-aerated soil, they are functionally important in arid-region, salt-affected soils. Algae are eukaryotic cells ranging in size from 2 to 20 mm. Several hundred algal species have been identified in soil, but only a small number of species are prominent throughout the world. Most algae grow best under moist to wet conditions, but some thrive in hot or cold environments.

DETERMINATION OF SOIL MICROBIAL COMMUNITIES A number of techniques have been developed to culture, identify, and characterize microbial communities. Such methods characterize substrate utilization by in situ microbial communities or by part of the community that is culturable in Biolog GN microplates, which are used to produce community-level physiological profiles. As no single universal method is currently available to determine the whole microbial community, a combination of methods was developed over time. Until recently, methods to analyze microbial 1071

1072

Microbial Communities

Table 1 Relative numbers and biomass of microbial species at 0–15-cm depth of soil Microorganisms

Number per gram of soil

Biomass (g/m2)

Bacteria

108–109

40–500

Actinomycetes

107–108

40–500

5

Fungi

6

10 –10

100–1500

Algae

104–105

1–50

[11]

(From Ref.

.)

community structure and function have been generalized for dominant groups such as bacteria and fungi. Population may be determined through direct counting,[12,13] specific cell component analysis,[14–16] substrate-induced respiration,[17] chloroform fumigation and microwave incubation, and extraction methods[18–20] and DNA analysis.[21] The use of molecular technology, including polymerase chain reaction amplification, cloning, and sequence analysis of the 16S subunit of rRNA, is becoming a much more common means of estimating the composition of the bacterial community.[21] FACTORS AFFECTING SOIL MICROBIAL COMMUNITIES The soil ecosystem comprises a variety of microhabitats with different physicochemical gradients and discontinuous environmental conditions that are influenced by the physical and chemical properties of the soil, the type of vegetative cover, the climate,[22] and the soil-crop management practices. The structure and function of soil microbial communities reflect between hosts of biotic and abiotic components of the terrestrial ecosystem. Amount, Quality, and Placement of Organic Residues Organic C compounds, from plant residues and soil organic matter, are used as energy and food sources by the heterotrophic microorganisms in soil. The quality of the plant litter reflects the biochemical composition of the substrates and the physical availability of the substrates to the microorganisms.[23] Bacteria tend to respond most rapidly to additions of simple C compounds such as starch, sugars, and amino acids, while fungi and actinomycetes dominate if complex C compounds such as cellulose and more resistant lignin materials are available.[24] In addition, if organic residues are deposited on the soil surface (under conifer forest and no-till systems), fungi dominate the microbial activity. Bacteria commonly play a larger


role if the substrates are mixed into the soil, as by earthworms, root distribution, and tillage.[25] Soil Moisture, Aeration, and Temperature While most microbes are aerobic and use oxygen as the electron acceptor in their metabolism, some microbes are anaerobic and use chemical compounds with combined oxygen other than gaseous oxygen. Facultative microbes use either aerobic or anaerobic forms of metabolism. Usually, these three metabolisms simultaneously occur in different habitats within a soil. Soil microbial activity is generally greatest when temperatures are within 20–40 C. The warmer end of this range tends to favor actinomycetes. Except for certain cryophilic species, most microorganisms cease metabolic activity below about 5 C, a temperature sometimes referred to as biological zero.[2,10] Soil Reaction Soil reaction is a factor in determining which specific microorganisms dominate and function in a particular soil. Although some bacterial species may thrive in any chemical condition of the soil, near-neutral pH results in the largest and most diverse composition of bacterial populations. Acidity enhances soil fungi activity. The acidity effect explains why fungi tend to dominate in forested soils while bacterial biomass generally dominates over fungi in slightly acidic subhumid to semiarid prairie and rangeland soils. Management Practices Type, timing, and degree of tillage affect microbial processes by placing plant residues on soil or in soil contact, exposing native and protected C, and by altering soil aeration, moisture, and temperature conditions.[26,27] Once in the soil, plant residues serve as energy and C sources for heterotrophic organisms.[28] In general, greater additions of plant residue to soil under conservation management practices are almost always followed by a flush of biological activities, coinciding with an increase in microbial biomass and metabolic activity.[27] The incorporation of easily decomposable plant residues with relatively low C : N ratio is a unique feature of legume-based cropping system, which enhances soil biological processes compared to conventional systems.

ROLE OF SOIL MICROBIAL COMMUNITIES The processes and functions of the terrestrial ecosystem are regulated by the activities of soil microorganisms.


Some of the important ecological functions related to the activities of the microbial community are litter decomposition, nutrient cycling, bioremediation, C sequestration, and improvement of the physical properties of the soil. Litter Decomposition and Nutrient Availability Litter decomposition is perhaps the most significant contribution of the soil microbes in terrestrial ecosystem. Through this process, the complex chemical compounds in dead leaves, roots, and other plant tissues are broken down into simpler chemical compounds, converting organically held nutrients into mineral forms available for renewed plant uptake. The release of N from decomposition of organic residues is a prime example. Soil microbes also assimilate wastes from animals and other organic materials added to soils. As a by-product of their metabolism, microbes synthesize new compounds, some of which help stabilize soil structure while others contribute to formation of soil organic matter and improve soil quality over time.[27] Inorganic Nutrient Transformations The transformation of inorganic nutrient elements and compounds is of great significance to the ecological functions of the soil systems, including plant growth. Nitrate, sulfate, and, to a lesser degree, phosphate ions are present in soil primarily because of microbial activities. Bacteria and fungi assimilate some of the N, P, and S in the organic materials they digest. Excess amounts of these nutrients may be excreted into the soil system in inorganic forms usable either by the bacteria and fungi themselves or by other organisms that feed on them. In this process, organically bound forms of N, P, and S are converted into inorganic forms that are available to higher plants. Likewise, the solubility and availability of the other essential elements, such as Fe and Mn, are largely determined by microbial activities in soil. In well-drained soils, these elements are oxidized by autotrophic microorganisms to their valence states, in which forms they become quite insoluble, which keeps Fe and Mn mostly in nontoxic forms even under fairly acidic conditions. Microbial oxidation also controls the potential for toxicity in soil contaminated with Se, As, and Cr compounds. Biological Nitrogen Fixation Nitrogen is one of the primary nutrients for plant growth. Atmospheric N2 is an inert gas that cannot


1073

be directly used by higher plants. Biological N fixation is one of the most important microbial processes in soils. Actinomycetes of the genus Frankia fix major amounts of atmospheric N in forest ecosystems; cyanobateria are important in flooded rice systems, wetlands, and deserts; and rhizobia are the most important group for the fixation of gaseous inert N in agricultural soils.[10] By far, the greatest amount of N fixation by these microorganisms occurs in root nodules or in other close associations with plants. Worldwide, every year, microbes in soils fix enormous quantities of atmospheric N into forms usable by higher plants.[2] Soil Aggregate Stability and Carbon Sequestration Aggregate structure is a key biologically mediated soil property that results from close associations of fungal hyphae with plant roots and from release of low molecular weight organic acids associated with clays, metals, and microaggregates.[29] In addition to C retention through efficient cell assimilation, extensive fungal hyphal associations with plant roots can contribute to protect and accumulate C by enhancing macroaggregation through physical enmeshing of soil microaggregates or releasing extracellular polysaccharide as binding agents or stabilization through complex association with metals as microaggregates.[30,31] The microbially derived decomposition products cement primary particles, organic debris, and microaggregates together to form and stabilize macroaggregates.[32] An enhanced physical protection of fragmented organic debris and microbial decomposition products within macroaggregates is an important mechanism enabling the soil C accumulation.[32] These mechanisms, in turn, increase the proportion of C in soil aggregates that is physically protected from microbial decomposition and protects soil from accelerated soil erosion.[27] Breakdown of Toxic and Xenophobic Compounds Many organic compounds toxic to plants or animals find their way into the soil. Some of these toxins are produced by soil microbes as metabolic by-products; some are applied as agrochemicals to kill pests and insects and to control weeds; and others are deposited in the soil because of environmental contamination. If these compounds continuously accumulate, they would do enormous ecological damage. Some toxins are xenophobic compounds foreign to biological systems, and these may resist attack by commonly occurring microbes. Soil bacteria and fungi are especially important in helping maintain a nontoxic soil environment by breaking down toxic compounds.

1074


Plant Protection and Diseases

Competition Interactions

Certain soil microbes protect higher plants from invasion by soil parasites and pathogens; however, others attack plant roots. According to several recent research studies, many plants that are resistant to soilborne diseases have camouflaged rhizosphere;[10,33] that is, although the rhizosphere of these resistant plants is enriched in microbial numbers, the types of microbes are more similar to those in the bulk soil than in the rhizosphere of disease-susceptible plants. Having a rhizosphere microbial community that is similar to the bulk soil community may make the rhizosphere of these plants less identifiable to pathogens. Beneficial rhizobacteria have an intriguing mode of action called induced systematic resistance, which prevents infection of plants by diseases, insects, or pests, both above and below ground. In many cases studied so far on crops, the resistance-inducing organism was a Pseudomonas or a Serratia bacterium.[10,33] This mechanism was shown to effectively reduce damages by numerous fungal, bacterial, and viral pathogens, and several leaf-eating insect pests.[33] Disease infestations occur in great variety and are induced by many different soil microorganisms. Among the microorganisms, bacteria, actinomycetes, and fungi are responsible for most of the common soilborne diseases of crops. Examples of such diseases are wilts, damping-off, root rots, clubroot of cabbage, and blight of potatoes.

Competition occurs when the growth of one species or community is suppressed by another species or community that more effectively obtains a limiting resource, or when one species or community is inhibited by the products of a second species or community. Limiting resources may be electron acceptors,[37,38] mineral resources,[39] host organisms,[40] and carbon sources.[41] When fresh organic residues are added to the soil, the vigorous heterotrophic soil microbes compete with each other for this food resource. If simple compounds such as sugars, starches, and amino acids are available in the organic residues, the bacteria initially dominate because of their rapid growth and preference for simple compounds. As these simple compounds are metabolized, the fungi, and particularly the actinomycetes, become more competitive in soil by altering habitat conditions. Certain microbes alter the acidity level in their vicinity to the disadvantage of competitors. Others produce strong bonding compounds to protect the food source. Production of antimicrobial compounds may also limit the development of other microbes. The potato soft-rot pathogen, Erwinia carotovora subsp. atroseptica (van Hall) Dye, is controlled by the production of the antibiotic 2,4-diacetylphloroglucinol by Pseudomonas fluorescens (Trevisan) Migula F113.[42] However, the production of antibiotics does not eliminate other possible mechanisms of action. It has been reported that nonpathogenic strains of Agrobacterium prevent expression of disease rather than killing the pathogens.[43]

INTERACTIONS OF SOIL MICROBIAL COMMUNITIES With the diversity in a soil ecosystem, many types of interactions of the microbial community with plant roots and other soil organisms, and between several microorganisms, are reported. The most important associations are given below. Antagonistic Interactions Among the antagonistic interactions, parasitism, where one species resides within another, and predation, where one species actively pursues another, are deleterious associations of two microbial populations in the sense that one organism gains an advantage at the expense of another. Bacteria that prey on other bacteria are active in soil ecosystems.[2,34,35] Protozoa and nematodes[36] that feed on bacteria are also active in the soil ecosystem. This can have a significant effect on population.[37] Parasitism of Pythium coloratum, the fungi responsible for cavity-spot disease in carrots, by actinomycetes has been reported.[38]


Mutualistic Interactions A mutually beneficial relationship between and among microorganisms is common in soil. The most familiar is the relationship between leguminous plants and rhizobia to fix atmospheric N. Symbiotic relationships are also found between populations of N-fixing communities and nonfixing communities,[44] and Rhizobium and arbuscular fungi.[45] One of the most ecologically and economically important activities of soil fungi is the mutually beneficial association between certain fungi and the roots of higher plants. This association is called mycorrhizae, a term meaning ‘‘fungus root.’’ In natural ecosystems, especially in degraded sites, many plants are quite dependent on mycorrhizal relationships and cannot survive without them; however, both parties benefit from such relationship. The microbiotic crusts are intricate living systems in arid rangelands. They consist of mutualistic associations that usually include algae or cyanobacteria along


with fungi, mosses, bacteria, and liverworts.[10] The presence of microbiotic crusts is considered a sign of a healthy or recovering ecosystems.

CONCLUSIONS The soil is a complex and dynamic component of the terrestrial ecosystem with a diverse community of microbes dominated by bacteria, fungi, actinomycetes, and algae. The microbial community is vital to the cycle of life on earth. They incorporate plant and residues into the soil and digest them, returning CO2 to the atmosphere, where it can be recycled through higher plants. Simultaneously, they release nutrients during decomposition of organic residues and form soil organic matter, which is vital for soil quality. With the diversity in a soil ecosystem, many types of community interactions are possible, including parasitism, competition, and symbiosis.

REFERENCES 1. Whipps, J.M. Microbial interactions and biocontrol in the rhizosphere. J. Exp. Bot. 2001, 52, 487–511. 2. Tate, R.L., III. Soil Microbiology; John Wiley & Sons, Inc.: New York, 1995; 147–170. 3. Muller, A.K.; Westergaard, K.; Christensen, S.; Sorensen, S.J. The diversity and function of soil microbial communities exposed to different disturbances. Microb. Ecol. 2002, 44, 49–58. 4. Naeem, S.; Thompson, L.J.; Lawler, S.P.; Lawton, J.H.; Woodfin, R.M. Empirical evidence that declining species diversity may alter the performance of terrestrial ecosystems. Philos. Trans. R. Soc. Lond. 1995, 347, 249–262. 5. Tilaman, D.; Wedin, D.; Knops, J. Productivity and sustainability influenced by biodiversity in grassland ecosystems. Nature 1996, 379, 718–720. 6. Hopper, D.U.; Vitrousek, P.M. The effects of plant composition on ecosystems processes. Science 1997, 277, 1302–1305. 7. Tilman, D.; Knops, J.; Wedin, D.; Reich, P.; Ritche, M.; Siemann, E. The influence of functional diversity and composition on ecosystem processes. Science 1997, 277, 1300–1302. 8. Wardle, D.A.; Zackrisson, O.; Horberg, G.; Gallet, C. The influence of island area on ecosystem properties. Science 1997, 277, 1296–1299. 9. Rosello-Mora, R.; Amann, R. The species concept of prokaryotes. FEMS Microbiol. Rev. 2001, 25, 39–67. 10. Brady, N.C.; Weil, R.R. Organisms and ecology of the soil. In Nature and Properties of Soils, 13th Ed.; Prentice Hall: New Jersey, 2002; 449–497. 11. Adu, J.K.; Oades, J.M. Utilization of organic materials in soil aggregates by bacteria and fungi. Soil Biol. Biochem. 1978, 10, 117–122.


1075

12. Alexander, M. Most probable number method for microbial populations. In Methods of Soil Analysis, Part II. Chemical and Microbiological Properties, 2nd Ed.; Page, A.L., Miller, R.H., Keeney, D.R., Eds.; American Society of Agronomy: Madison, WI, 1982; Vol. 9, 815–820. 13. Bottomley, P.J. Light microscopic methods for studying soil microorganisms. In Methods of Soil Analysis, Part II. Microbiological and Biochemical Properties; Weaver, R.W., Angle, J.S., Botomley, P.S., Eds.; Soil Science Society of America, Inc.: Madison, WI, 1994; 81–106. 14. Webster, J.J.; Hampton, G.J.; Leach, F.R. ATP in soil: a new extractant and extraction procedure. Soil Biol. Biochem. 1984, 16, 335–342. 15. Joergensen, R.G. Quantification of the microbial biomass by determining ninhydrin-reactive N. Soil Biol. Biochem. 1996, 28, 301–306. 16. White, D.; Davis, W.M.; Nickels, J.S.; King, J.D. Determination of the sedimentary microbial biomass by extractable lipid phosphate. Oecologia 1979, 40, 51–62. 17. Anderson, J.P.E.; Domsch, K.H. A physiology method for quantitative measurement of microbial biomass in soils. Soil Biol. Biochem. 1978, 10, 519–525. 18. Jenkinson, D.S.; Powlson, D.S. The effects of biocidal treatments on metabolism in soil. V. A method for measuring soil biomass. Soil Biol. Biochem. 1976, 8, 209–213. 19. Vance, E.D.; Brooks, P.C.; Jenkinson, D.S. An extraction method for measuring soil microbial biomass-C. Soil Biol. Biochem. 1987, 19, 703–707. 20. Islam, K.R.; Weil, R.R.; Mulchi, C.L.; Glenn, S.D. Freeze-dried soil extraction method for the measurement of microbial biomass C. Biol. Fertil. Soils 1997, 24, 205–210. 21. Marilley, L.; Vogt, G.; Blanc, M.; Aragno, M. Bacterial diversity in the bulk soil and rhizosphere fractions of Lolium perenne and Trifolium repens as revealed by PCR restriction analysis of 16S rDNA. Plant Soil 1998, 198, 219–224. 22. Clark, F.E.; Paul, E.A. The microflora of grassland. Adv. Agron. 1970, 22, 373–435. 23. Wardle, D.A.; Giller, K.E. The quest for a contemporary ecological dimension to soil biology. Soil Biol. Biochem. 1996, 28, 1549–1554. 24. Moller, J.; Miller, M.; Kjoller, A. Fungal–bacterial interactions on beech leaves: influence on decomposition and dissolved organic C quality. Soil Biol. Biochem. 1999, 31, 367–374. 25. Beare, M.H.; Parmelee, R.W.; Hendrix, P.F.; Cheng, W.; Coleman, D.C.; Crossley, D.A. Microbial and faunal interactions and effects on litter N and decomposition in agroecosystems. Ecol. Monogr. 1992, 6, 569–591. 26. Dick, W.A. Tillage system impacts on environmental quality and soil biological parameters. Soil Tillage Res. 1997, 41, 165–167. 27. Islam, K.R.; Weil, R.R. Soil quality indicator properties in mid-Atlantic soils as influenced by conservation management. J. Soil Water Conserv. 2000, 55, 69–78. 28. Van de Geijn, S.C.; Van Veen, J.A. Implication of increased carbon dioxide levels for carbon input and turnover in soils. Vegetation 1993, 104=105, 283–292.

1076

29. Tisdall, J.M.; Oades, J.M. Organic matter and waterstable aggregates in soils. J. Soil Sci. 1982, 33, 141–163. 30. Oades, J.M. Soil organic matter and structural stability: mechanisms and implications for management. Plant Soil. 1984, 76, 319–337. 31. Chenu, C. Influence of a fungi polysaccharide, scleroglucan, on clay microstructures. Soil Biol. Biochem. 1989, 21, 299–305. 32. Jastrow, J.D.; Miller, R. Soil Aggregation in the Rhizosphere: Optimal Conditions for Multiple Mechanisms; The Ecol. Soc. Amer. 85th Ann. Meeting, Snowbird, UT, 2000; 21 pp. 33. Imran, A.S.; Shaukat, S.S. Mixtures of plant disease suppressive bacteria enhance biological control of multiple tomato pathogens. Biol. Fertil. Soils 2002, 36, 260–268. 34. Casida, L.E. Competitive ability and survival in soil of Pseudomonas strain 679-2, a dominant, nonobligate bacterial predator of bacteria. Appl. Environ. Microbiol. 1992, 58, 32–37. 35. Germida, J.J.; Casida, L.E. Ensifer adherens predatory activity against other bacteria in soil, as monitored by indirect phage analysis. Appl. Environ. Microbiol. 1983, 45, 1380–1388. 36. Mattison, R.G.; Harayama, S. The predatory soil flagellate Heteromita globosa stimulates toluene biodegradation by a Pseudomonas sp. FEMS Microbiol. Lett. 2001, 194, 39–45. 37. Mikola, J.; Sulkava, P. Responses of microbial-feeding nematodes to organic matter distribution and predation in experimental soil habitat. Soil Biol. Biochem. 2001, 33, 811–817. 38. El-Tarabily, K.A.; Hardy, G.E.; Sivasithamparam, K.; Hussein, A.M.; Kurtbo¨ke, D.I. The potential for the biological control of cavity-spot disease of carrots,



39.

40.

41.

42.

43.

44.

45.

caused by Pythium coloratum, by streptomycete and non-streptomycete actinomycetes. New Phytol. 1997, 137, 495–507. Kotsyurbenko, O.R.; Glagolev, M.V.; Nozhevnikova, A.N.; Conrad, R. Competition between homoacetogenic bacteria and methanogenic archaea for hydrogen at low temperature. FEMS Microbiol. Ecol. 2001, 38, 153–159. van Bodegom, P.; Stams, F.; Mollema, L.; Boeke, S.; Leffelaar, P. CH4 oxidation and the competition for O2 in the rice rhizosphere. Appl. Environ. Microbiol. 2001, 67, 3586–3597. O’Hara, G.W. Nutritional constraints on root nodule bacteria affecting symbiotic nitrogen fixation: a review. Aust. J. Exp. Agric. 2001, 41, 417–433. Cronin, D.; Moenne-Loccoz, Y.; Fenton, A.; Dunne, C.; Dowling, D.N.; O’Gara, F. Ecological interaction of a biocontrol Pseudomonas fluorescens strain producing 2,4-diacetylphloroglucinol with the soft rot potato pathogen Erwinia carotovora subsp. atrosetica. FEMS Microbiol. Ecol. 1997, 23, 95–106. Johnson, K.B.; DiLeone, J.A. Effect of antibiosis on antagonist dose–plant disease response relationships for the biological control of crown gall of tomato and cherry. Phytopathology 1999, 89, 974–980. Bever, J.D.; Simms, E.L. Evolution of nitrogen fixation in spatially structured populations of Rhizobium. Heredity 2000, 85, 366–372. Biro, B.; Koves-Pechy, K.; Voros, I.; Takacs, T.; Eggenberger, P.; Strasser, R.J. Interrelations between Azospirillum and Rhizobium nitrogen-fixers and arbuscular mycorrhizal fungi in the rhizosphere of alfalfa in sterile, AMF-free or normal soil conditions. Appl. Soil Ecol. 2000, 15, 159–168.

Micromorphology and Soil Quality Ahmet R. Mermut University of Saskatchewan, Saskatoon, Saskatchewan, Canada

INTRODUCTION The book Micropedology[1] and a large number of publications have induced many people to practice soil micromorphology since 1950s. Major areas where micromorphological analyses were used to determine soil quality in the 1970s and 1980s were in soil physics (soil pores, aggregates, and crusting), biology including ecosystems and faunal activity, soil chemistry, environmental chemistry, and mineralogy. These applications were made possible by the advent in computer science and development of image analyses technologies[2,3] for better quantification of soil features. So far, quantitative data were obtained from twodimensional measurements, on the basis of aerial percentage. As is performed in stereology, two-dimensional measurements are extended to volume percent by the Delesse principle; that is, aerial percentage is an unbiased estimate of volume percentage.[4]

SOIL PORES, AGGREGATES, TILLAGE, AND SOIL CRUSTING Of all the physical measurements, pore size distribution is the most pertinent parameter for plant growth.[5] Soil porosity is dynamic[6] and exhibits a high degree of physical anisotropy. Primary focus in porosity studies has been on the measurement of total porosity and shape and size classes of pores. It was discovered that measurements of porosity itself are not sufficient unless it is coupled with other soil physical, biological, and chemical analyses.[7] Despite the difficulties, attempts were made to measure the full spectrum of pore sizes, from pore diameters in the order of centimeters to those of nanometers.[8] Application of micromorphology in soil structure and aggregates gained considerable momentum in the 1960s, 1970s, and 1980s (Fig. 1).[9–12] Impact of tillage implements, including zero tillage and crop rotation on soil structure, received considerable attention.[13] This was followed by studies dealing with surface crusting and sealing.[14–16] These studies suggested that the susceptibility of soils to surface crusting and sealing Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006651 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

depends on a combination of several physical and chemical properties of soils (Fig. 2). Several investigators have studied the effect of different kinds of tillage operations on soil porosity and structure (Table 1).[17–19] The data led to insight of different properties and functions of soil structure and porosity. Pagliai[20] suggested that conservation or zero-tillage practices, when compared with conventional tillage, maintain soil structure and optimize soil porosity. Several workers have attempted to directly quantify porosity using thin sections or impregnated and polished soil blocks.[21–23] Yet the influence of tillage operations on soil porosity is not completely appreciated. Mermut, Grevers, and de Jong[24] used image analyses to evaluate the impact of deep ripping and paraplowing on soil porosity in two lacustrine soils in Saskatchewan. Deep ripping increased soil porosity in the surface horizon, whereas the influence of paraplowing was evident up to 30 to 60 cm depth (Fig. 3). Two tillage methods used in this study increase large planar voids, but nonplanar intra-aggregate voids were not influenced by tillage. It has been suggested that pores should be measured more frequently and over a long time, so that a clearer picture of the dynamics is better understood. Mineralogy of clay sized particles and rainstorm characteristics are among the major factors that determine the nature of soil sealing.[25] Southard, Shainberg, and Singer[26] found that soil material suspended in water containing electrolytes resulted in less dense and more porous crust. Several crust types were identified.[15,16,27]

SOIL FAUNA The impact of organisms on the function and quality of soil is an important area of research.[28,29] By using interactive image analysis system and a video camera, micromorphologists can measure the size and shape of fecal pellets of animals observed in soil thin sections. Such an analysis can determine the possible ecological significance of the fecal pellets of animals in soils, and the data generated are essential to the protection, use, and management of soils and forests (Fig. 4). 1077

1078

Micromorphology and Soil Quality

Fig. 1 (A) Thin section micrograph showing interconnected simple and compound, and (B) packing void (interaggregate pores) under partially crossed polarized light (p ¼ pores; q ¼ quartz: c ¼ soil aggregates). Right: high-resolution TEM micrographs of an ultracut of primary aggregate (p ¼ porosity; k ¼ kaolinite; h, I, and x are various soil minerals). (From Ref.[8].)

Table 1 Average porosity of the upper 1-cm thick layer of soil measured by image analyses after different irrigation system

Sprinkler

Furrow irrigation track

Furrow irrigation ridge

Bare soil

16

5

16

Cropped soil

14

16

17

Seed bed 29

Values are expressed in % of porosity. (From Ref.[19].)

SOIL ORGANIC MATTER Micromorphologists are also interested in soil organic matter and its dynamics.[30] Several textbooks and articles were published by Kubiena[31] between 1938 and 1970 on soil organic matter, mainly to determine the function and turnover of organic matter in soils. Submicroscopic techniques were also used to study

Fig. 2 Scanning electron micrograph of a crust. Note the fine particles covering the surface top central part (the white bar is 100 mm). (From Ref.[15].)


organic matter in situ by histochemical staining methods in terms of organic matter function in ecosystem.[32] This is an important and high-priority area of research.

MICROCHEMISTRY AND SUBMICROSCOPIC STUDIES Pioneering works in undisturbed in situ microchemical analyses of polluted soils have been carried out on heavy metals.[33] When coupled with image analyses, these techniques provide not only the location, but also quantity and distribution of the pollutant in soils. With the help of the new instrumentation, such as synchrotron, revolutionary scientific research can be performed.

Fig. 3 Contact images of this section representing the control, deep ripping, and paraplow parcel from the Tisdale site, Saskatchewan, Canada. (From Ref.[23].) (View this art in color at www.dekker.com.)


Fig. 4 (A) A soil thin section showing the fecal pellets of a soil-ingesting animal in the surface soil, and (B) more rounded fecal pellets. The frame length is about 0.5 cm.

CONCLUSIONS Soil micromorphology has numerous applications to soil quality assessment. These include assessment of soil pores, aggregates, crusting, biology including ecosystems and faunal activity, soil chemistry, environmental chemistry, and mineralogy. Whereas numerous advances were made in the 1970s and 1980s, there is a vast scope for future use of these techniques.

REFERENCES 1. Kubiena, W. Micropedology; Collegiate Press: Ames, IA, 1938. 2. Miedema, R.; Mermut, A.R. Soil micromorphology. In An Annotated Bibliography 1968–1986; CAB International: Wallingford, Oxen, UK, 1990. 3. Protz, R.; Sweeney, S.J. An application of spectral image analyses to soil micromorphology. 1. Methods of analyses. Geoderma 1992, 53, 275–287. 4. Weibel, E.R. Stereological Methods; Academic Press: London, U.K., 1979; Vol. 1. 5. Thomasson, A.J. Towards an objective classification of soil structure. J. Soil Sci. 1978, 29, 38–46. 6. Cassel, D.K. Spatial and temporal variability of soil physical properties following tillage of norfolk loamy sand. Soil Sci. Soc. Am. J. 1983, 47, 196–201. 7. Bouma, J.; Jongreius, A.; Boersma, O.; Jager, A.; Schoonderbeck, D. The functions of different types of macro-pores during saturated flow through four swelling soil horizons. Soil Sci. Soc. Am. Proc. 1977, 41, 945–950. 8. Bui, E.; Mermut, A.R.; Santos, M.C.D. Microscopic and ultramicroscopic porosity of an oxisol form pernambuco, Brazil. Soil Sci. Am. J. 1989, 53, 661–665. 9. Bouma, J. Microstructure and Stability of Two Sandy Loam Soils with Different Soil Management; Agricultural Research Reports; PUDOC: Wageningen, The Netherlands, 1969; Vol. 724. 10. Tokaj, J. Micromorphological investigations on soil aggregates. Third international working meeting on soil micromorphology, wroclaw, poland. Zesz. Probl. Postep. Nauk Rol. 1972, 123, 733–746.


1079

11. de Meeter, T.; Jongerius, P.D. The relationship between soil erodability factor K (universal soil loss equation, aggregate stability and micromorphological properties of soils in the Hornos area, Southern Spain. Earth Surf. Processes 1978, 3, 379–391. 12. Santos, M.C.D.; Mermut, A.R.; Ribeiro, M.R. Submicroscopy of clay aggregates in an oxisol from pernambuco, Brazil. Soil Sci. Soc. Am. J. 1989, 53, 575–582. 13. Jongerius, A. Some micromorphological aspects of regrouping phenomena in Dutch soil. Geoderma 1970, 4, 311–331. 14. Pagliai, M.; Bisdom, E.B.A.; Ledin, S. Changes in surface structure (crusting) after application of sewage sludge and pig slurry to cultivated agricultural soil in Northern Italy. Geoderma 1983, 30, 35–53. 15. Arshad, M.A.; Mermut, A.R. Micromorphological and physico-chemical characteristics of soil crust types in Northwestern Alberta, Canada. Soil Sci. Soc. Am. J. 1988, 52, 724–729. 16. Bresson, L.M.; Valenti, C. Comparative Micromorphological Studies of Soil Crusting in Temperate and Arid Environments, Transactions 14th International Congress of Soil Science, Kyoto, Japan, August, 12–18, 1990; 236–243. 17. Pawluk, S. Micromorphological investigations of cultivated Gray Luvisols under different management practices. Can. J. Soil Sci. 1980, 60, 731–745. 18. Collins, J.F.; Larney, F. Micromorphological observations of compacted horizons (cultivated pans) from various horizons in Irish Tillage Soils. In Soil Micromorphology; Fedoroff, N., Bresson, L., Courty, M.A., Eds.; Association Francaise, 1987; 451–457. 19. Tedeschi, A.; Melle, G.; Terribile, F. Soil pore geometry changes in surface horizons, under different irrigation regime. 17th World Congress of Soil Science; Transactions International Union of Soil Science (IUSS); Bangkok, Thailand, 2002; Symposium 35, 1–12, 264 pp. 20. Pagliai, M. Effects of different management practices on soil structure and surface crusting. In Soil Micromorphology; Fedoroff, N., Bresson, L., Courty, M.A., Eds.; Association Francais pour l’Etude du Sol, 1987; 415–421. 21. Shipitalo, M.J.; Protz, R. Comparison of morphology and porosity of a soil under conventional and zero tillage. Can. J. Soil Sci. 1987, 67, 445–456. 22. Moran, C.J.; McBratney, A.B.; Koppi, A.J. A rapid method for analyses of soil micropore structure. I. Specimen preparation and digital binary image production. Soil Sci. Soc. Am. J. 1989, 53, 921–928. 23. Grevers, M.C.J.; de Jong, E. Soil structure changes in sunsoiled solonetzic and chernozemic soils measured by image analyses. In Digitization, Processing and Quantitative Interpretation of Image Analysis; Geoderma; Mermut, A.R., Norton, D.L., Eds.; Elsevier: Amsterdam, 1992; Vol. 53, 289–307. 24. Mermut, A.R.; Grevers, M.C.J.; de Jong, E. Evaluation of pores under different management systems by image analysis of clay soils in saskatchewan, Canada. In Digitization, Processing and Quantitative Interpretation of Image Analysis; Geoderma; Mermut, A.R., Norton, D.L., Eds.; Elsevier: Amsterdam, 1992; Vol. 53, 357–372.

1080

25. Mermut, A.R.; Luk, S.H.; Romkens, M.J.M.; Poesen, J.W.A. Micromorphological and mineralogical components of surface sealing in loess soils from different geographic regions. Geoderma 1995, 66, 71–84. 26. Southard, R.J.; Shainberg, I.; Singer, M.J. Effect of electrolyte concentration on the micromorphology of artificial depositional crust. Soil Sci. 1988, 145, 278–288. 27. Onofiok, O.; Singer, M.J. Scanning electron microscope studies of surface crusts formed by simulated rainfall. Soil Sci. Soc. Am. J. 1984, 48, 1137–1143. 28. Babel, U. Echytraeen-losungsgefuge in Loess (Dropping fabrics from Anchytraeidae in Loess). Gedoderma 1969, 2, 57–63. 29. Kositra, M.J. The interpretation and classification of features produced by pelecypods (Mollusca) in marine



30.

31. 32. 33.

intertidal deposits in the Netherlands. Geoderma 1981, 26, 83–94. Bal, L. Micromorphological analyses of soils. In Lower Level in the Organization of Soil Organic Materials; Soil Survey Papers; Netherlands Soil Survey Institute: Wageningen, 1973; Vol. 6, 174 pp. Kubiena, W. Micromorphological Features of Soil Geography; Rutgers University: New Brunswick, 1970. Foster, R. In situ localization of organic matter in soils. Quaest. Entomol. 1985, 21, 609–633. Bisdom, E.B.A.; Henstra, S.; Werner, H.W.; Boudewijn, P.R.; Knippenberg, W.F.; deGrefte, H.A.M.; Gourgout, J.M.; Migeon, H.N. Quantitative analyses of trace and major elements in thin sections of soils with the secondary ion microscope (Cameca). Geoderma 1983, 30, 117–134.

Mineland Reclamation and Soil Carbon Sequestration Vasant A. Akala Rattan Lal The Ohio State University, Columbus, Ohio, U.S.A.

INTRODUCTION Revegetation of mined lands is one of the management options for mitigation of the negative impacts of mining. The area of minelands that have been reclaimed since 1970 is 1 Mha in the United States alone.[1–2] Land restoration measures can reverse the degradation trends, leading to positive changes in the ecosystem. These include improved water, air, and soil qualities and the attendant socio-economic benefits.[3] An important ancillary benefit of mineland reclamation is its potential to sequester soil organic carbon (SOC).[4] The potential for sequestration of carbon (C) in the aboveground biomass and minesoils of reclaimed minelands may be high and merit serious consideration as a significant benefit of restoration of highly disturbed land.

The potential of SOC enhancement in minelands depends on biomass productivity, root development in sub-soil, and changes in minesoil properties resulting from overburden weathering.[11] Improvement in soil aggregation is an important factor influencing minesoil development.[12,13] The changes in minesoil properties impact on biomass productivity and pedogenic processes that lead to SOC sequestration in minesoils. Soil processes that are important in the C dynamics of minesoils and those that lead to SOC sequestration are weathering, stable structure formation, erosion, compaction, aeration, aggregation, nutrient recycling, humification, and mineralization. Table 1 shows how some of these processes relate to SOC sequestration in minesoils.

CASE STUDY MINESOIL DEVELOPMENT AND SOIL ORGANIC CARBON SEQUESTRATION The rate of formation of minesoil or minesoil development over time may be the single most important edaphic factor determining the success of the reclamation plan. The rate of minesoil development is a function of reclamation plan management. Freshly exposed mine spoil may be thought of as soil at time zero, in respect to formation, because it consists only of rock and pulverized rock material.[5] Mine spoils reflect the properties of the parent material more closely than natural soils because of the initial stage of pedogenic development.[6] Although there are different theories regarding the rate of minesoil development, there is a general consensus that soil weathering is rapid in the early stages of reclamation, and the rate of change decreases over time.[7–9] Very rapid chemical reactions occur in spoil as the exposed material begins to come into equilibrium with its new environment.[7] Also, minesoils are exposed to processes of physical weathering such as cycles of wetting and drying, freezing and thawing, and mechanical disruption by roots.[9] Rapid changes in particle size distribution take place as spoil material weathers.[10] Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001626 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

Akala and Lal[14] conducted a study on a chronosequence of reclaimed minelands, where the effects of reclamation duration (time), land use (pasture and forest, with and without topsoil application), and soil processes (aggregation and humification) on SOC sequestration were observed. The study showed that soil disturbance by mining resulted in high loss of the antecedent SOC pools and reclamation of mined lands regained SOC. Temporal changes in the SOC pool of reclaimed minesoils showed that the potential to sequester SOC was high. The SOC pool of the soil surface (0–15 cm depth) during the initial phases of reclamation was 10–15 Mg ha 1, which increased to 45–55 Mg ha 1 in 25–30 years (Fig. 1). The low SOC pool during the initial phases of reclamation was attributed to the drastic disturbance caused by mining. Soil that was removed and stored, and applied during reclamation lost significant amounts of SOC by decomposition and erosion during mining. Reclamation, minimum perturbation, and management of restored ecosystems, over a period of 20–25 years, had an ameliorative effect on minesoils because of decreases in soil compaction (reduction in soil bulk density), improvement in soil structure (high aggregation), and increases in SOC accumulation (enhanced biomass productivity). The SOC pool for 15–30 cm depth also increased over 1081

1082

Mineland Reclamation and Soil Carbon Sequestration

Table 1 Minesoil processes and soil organic carbon (SOC) sequestration Processes

Relation to SOC sequestration

Weathering

Formation of soils from the geologic material in the soil reduces the particle size of rock and coarse fragments, leading to the formation of stable peds and aggregates.

Erosion

Probably the most important process to be addressed in the initial stages of reclamation because the topsoil is susceptible to erosion. Control of erosion in the initial phases assures establishment of vegetation and favorable impact on environmental quality.

Aggregation

Increasing clay and organic matter content favor the eventual formation of minesoil peds and minesoil structure. The presence and development of minesoil structure by aggregation leads to SOC sequestration. Indirectly, aggregation reduces erodibility and increases pore space, thereby increasing infiltration and permeability, and thus improving minesoil quality.

Compaction

Vehicular traffic involved in implementing the reclamation plan causes compaction. This result in increased density and soil strength and decreased aeration of the re-graded spoil material. As a result of these changes, permeability is reduced and root proliferation and rooting depth are reduced. Water erosion may also be accelerated.

the period of reclamation. Possible mechanisms of C sequestration in reclaimed mineland were development of an A horizon, increased aggregation through formation of organomineral complexes, and humification of soil organic matter (SOM). Minesoil aggregation and C sequestration are closely related. Akala and Lal[14] observed a high correlation between water stable aggregates and SOC content. Land disturbance caused by mining not only decreased SOC content of the soil aggregates (secondary organomineral complexes) but also caused a drastic decrease in SOC content associated with the primary soil particles (primary organomineral complexes). The loss in SOC content (in comparison to the antecedent levels in control sites) was 65% in the clay fraction, 75% in the silt fraction, and 40% in the sand fraction. The temporal increase in SOC content in all particle size fractions over the reclamation duration reflected the

Fig. 1 Change in soil organic carbon (SOC) during reclamation over time.


increase in total SOC pool. The increase (difference in SOC contents between the end and beginning of reclamation period) was 3–6 times the antecedent level for the clay and silt fractions and 1–2 times for the sand fraction. A large increase in the clay fraction showed the reactive nature of minesoils and its ability to respond to changes in SOC pool. The reactive nature of the primary particles coupled with the increase in SOM input was the most important factors for SOC sequestration in the initial stages of mineland reclamation.

DISCUSSION The rates of SOC sequestration in reclaimed minelands can be 2–3 Mg ha1 yr1 for 0–15 cm depth and 1–2 Mg ha1 yr1 for 15–30 cm depth, and are higher than those observed in other land uses. For example, SOC sequestration rates by adoption of conservation tillage are in the range of 0.1–0.5 Mg ha1 yr1 in temperate regions and 0.05–0.2 Mg ha1 yr1 in tropical regions.[15] Similarly, SOC sequestration rates are in the range of 0.3–0.5 Mg ha1 yr1 by restoration of wetlands.[16] Rates of SOC sequestration reported above can be sustained for a period of 10–15 years. The rate of change of SOC pools or SOC sequestration in reclaimed minelands depends on the magnitude of loss by perturbation (mining) and how far the reclaimed sites are from attaining the equilibrium SOC level. Soils that initially have low SOC content and those that are converted to restorative land management have a potential SOC sink capacity. Reclaimed minelands are an example of such aggrading ecosystems. The SOC pool that can be potentially attained in reclaimed minelands in the United States is 25 million metric tonnes carbon.[15] The net emission of carbon

Mineland Reclamation and Soil Carbon Sequestration

dioxide (CO2) increased from 1037 Tg (1 Tg ¼ 1 Teragram ¼ 1 1012 g ¼ 1 million metric ton) C equivalent in 1990, to 1263 Tg C equivalent in 1996 in the United States.[17] Assuming that the reclaimed minelands were at least 25 years old, they had a potential to offset 4% of CO2 emissions for the United States. The importance of this potential may be realized when considered in combination with the restoration of all degraded soils and ecosystems and in the context of other soil management strategies.[15]

CONCLUSIONS The potential to sequester SOC in reclaimed minelands is high at 2–3 Mg ha1 yr1. SOC sequestration can be an environmentally friendly use of reclaimed minelands, and long-term C storage in such systems can offset part of the CO2 emissions (2–4%). The formation of primary organomineral complexes is the first step in the process of SOC sequestration. The total SOC content of the 25 to 30- year-old reclaimed sites can be 30–35% higher than that of the undisturbed sites, but not all of the SOC is sequestered in the secondary organomineral complexes, thereby making it vulnerable to future land disturbance. Soil aggregation and humification in reclaimed minelands can be higher than the undisturbed sites, resulting in high SOC sequestration. Pasture with topsoil application can achieve the highest SOC sequestration in a short period of time.

REFERENCES 1. Office of Surface Mining (OSM). 1998 Annual Report; Office of Surface Mining, Department of Interior: Washington, DC, 1998. 2. National Mining Association (NMA). Facts About Coal; National Mining Association: Washington, DC, 1998.


1083

3. Hossner, L.R., Ed.; Reclamation of Surface Mined Lands; CRC Press: Boca Raton, FL, 1991; Vols. I and II. 4. United States Department of Energy (USDOE). Energy Department Launches Thirteen New Research Projects to Capture and Store Greenhouse Gases; DOE Fossil Energy Techline: Washington, DC, 2000. 5. Kohnke, H. The reclamation of coal minesoils. Adv. Agron. 1950, 2, 317–349. 6. Sobek, A.A.; Smith, R.M.; Schuller, W.A.; Freeman, J.R. Overburden Properties that Influence Minesoils; NCA=BCR, 1976; 159 pp. 7. Struthers, P.H. Chemical weathering of strip minesoils. The Ohio J. of Sci. 1964, 64, 125–131. 8. Caspall, F.C. Soil Development on Surface Mine Spoils in Western Illinois; NCA=BCR, 1975; 228 pp. 9. Schafer, W.M.; Nielsen, G.A.; Dollhopf, D.J.; Temple, K. Soil Genesis, Hydrological Properties, Root Characteristics and Microbial Activity of 1 to 50 Year Old Strip Mine Spoils, EPA-600=7-79-100; USEPA: Cincinnati, OH, 1979. 10. Van Lear, D.H. Effects of Spoil Texture on Growth of K-31 Tall Fescue; Research Note NE-141; USDA Forest Service, 1971. 11. Haering, K.C.; Daniels, W.L.; Roberts, J.A. Changes in minesoil properties resulting from overburden weathering. J. Environ. Qual. 1993, 22 (1), 194–200. 12. Boerner, R.E.J.; Scherzer, A.J.; Brinkman, J.A. Spatial patterns of inorganic N, P availability, and organic C in relation to soil disturbances: a chronosequence analysis. Appld. Soil Ecology 1998, 7 (2), 59–177. 13. Malik, A.; Scullion, J. Soil development on restored opencast coal sites with particular reference to organic matter and aggregate stability. Soil Use Manage. 1998, 14 (4), 234–239. 14. Akala, V.A.; Lal, R. Potential of mineland reclamation for soil organic carbon sequestration in Ohio. Land Degrad. Develop. 2000, 11, 289–297. 15. Lal, R.; Kimble, J.M.; Follett, R.F.; Cole, C.V. The Potential of U.S. Cropland to Sequester Carbon and Mitigate the Greenhouse Effect; Ann Arbor Press: Chelsea, MI, 1998; 128 pp. 16. Mitsch, W.J.; Wu, X. Wetlands and global change. In Soil Management and Greenhouse Effect; Lal, R., Kimble, J., Levine, E., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1995; 205–230. 17. Environmental Protection Agency (EPA). Inventory of US Greenhouse Gas Emissions and Sinks: 1990–1996; Environmental Protection Agency: Washington, DC, 1998.

Minerals: Primary Nikolaos I. Barbayiannis Vissarion Z. Keramidas Soil Science Lab, Aristotle University of Thessaloniki, Thessaloniki, Greece

INTRODUCTION

Zircon

According to the Glossary of Soil Science Terms,[1] primary minerals are those that have not been altered chemically since their deposition or crystallization from molten lava (magma). Primary minerals identified in soils belong mainly to the classes of silicates, oxides of Fe, Zr, and Ti, and phosphates (apatite). Their study is essential because: a) they serve as sources of plant nutrients; b) they are, in certain cases, the precursors of secondary clay minerals; and c) they provide information about soil development.

Zircon is another nesosilicate mineral. Its structure consists of alternating edge-sharing silicon tetrahedra that are held together with Zr4þ ions that are located in the center of triangular dodecahedra.[3] Zircon is found as residual grains in the sand and silt fractions of soils and because of its stability in pedogenic environments it is frequently related to the degree of soil development and soil age. Pyroxenes and Amphiboles

SILICATES Of the primary minerals found in soils, silicates are the most abundant, comprising nearly 40% of the common minerals. The building unit of the silicates is the silicon tetrahedron. Silicate structures may consist of single tetrahedra (nesosilicates), double tetrahedra (sorosilicates), rings (cyclosilicates), single or double chains (inosilicates), sheets (phyllosilicates), or framework patterns (tectosilicates). Typical silicate minerals most likely to be found in soils are presented in Table 1.

Olivines Olivines are olive-green nesosilicates in which divalent cations mainly Mg2þ and Fe2þ join the silicon tetrahedra by forming octahedra of Mg2þ or Fe2þ. The forsterite–fayalite series include the most abundant naturally occurring olivines, with Mg2þ and Fe2þ being the respective cation of the series end members. Olivines are the first minerals to crystallize in the initial stages of magma solidification and they are highly unstable minerals because the high ratio of the divalent cations to silicon in its structure renders them vulnerable to chemical attack. Therefore, olivines are very easily weathered in soils and for this reason they are considered relatively rare soil constituents and can be found only in very young soils and in the coarser sand fractions. 1084 Copyright © 2006 by Taylor & Francis

Pyroxenes and amphiboles are ‘‘ferro-magnesian’’ minerals with single and double chain structures, respectively (Table 1). Pyroxene chains have silicon tetrahedra sharing two O atoms. In amphiboles the chains are formed by silicon tetrahedra sharing alternatively two and three O atoms. In both classes of minerals, the chains are held together mainly by divalent cations (Mg2þ, Fe2þ, or Ca2þ) located at the center of an octahedron.[4] Augite and hornblende are the most important minerals of the pyroxenes and amphiboles, respectively. Pyroxenes and amphiboles are formed at lower temperatures and pressures during magma solidification compared to olivines and are more stable in the weathering environment. In soils, however, they weather relatively faster and are largely confined in the sand and silt fractions. Some of them may occur in the clay fraction of soils that are young and have not been subjected to intensive weathering.[5] Feldspars Feldspars are tectosilicates, with three-dimensional framework of corner linked silicon and aluminum tetrahedra. The aluminum tetrahedral result from Al substitution for Si and the charge deficit created is compensated by Na, K, and Ca. Feldspars are virtually present in all sediments and soils in quantities that vary with the nature of the parent material and the stage of weathering. The majority of feldspars belong to the ternary system NaAlSi3O8 (albite) –KAlSi3O8 Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042713 Copyright # 2006 by Taylor & Francis. All rights reserved.

Minerals: Primary

1085

Table 1 Common primary minerals of the silicate class found in soils Silicon tetrahedra arrangement Nesosilicates (SiO4)

4

(single)

Name of mineral or group Olivine

Ideal formula (Mg,Fe)2SiO4

Zircon

ZrSiO4

Sorosilicates (Si2O7)6 (double)

Epidote

Ca2(Al,Fe)Al2O(SiO4)4(Si2O7)(OH)

Cyclosilicates (Si6O18)12 (rings) Inosilicates

Tourmaline

(Na,Ca)(Li,Mg,Al)(Al,Fe,Mn)6(BO3)3(Si6O18)(OH)4

(Single chains) (SiO3)2

Pyroxenes

(Double chains) (Si4O11)

6

Augite

(Ca,Na)(Mg,Fe,Al)(Si,Al)2O6

Hypersthene

(Mg,Fe)SiO3

Amphiboles Hornblende

2

Phyllosilicates (Si4O10)

(sheets)

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

Micas Muscovite

0

Tectosilicates (SiO2) (framework)

KAl2(AlSi3O10)(OH)2

Feldspars Orthoclase

KAlSi3O8

Albite

NaAlSi3O8

Anorthite

CaAl2Si2O8

Quartz

SiO2

(From Ref.[2].)

(K-feldspar) –CaAl2Si2O8 (anorthite). In K-feldspars one out of every four Si is replaced by Al, with K balancing the charge. Most common K-feldspars are sanidine, orhoclase, microcline, and adularia. Sodium and calcium feldspars form a series called ‘‘plagioclase’’ with compositions ranging from pure albite (Na-feldspar) to pure anorthite (Ca-feldspar). The presence of feldspars in soils is related to the overall mineralogical composition and the prevailing environmental conditions, including climate, topography, degree of leaching, presence of chelating agents, redox status, and soil solution composition.[6] Feldspars usually follow the stability sequence: anorthite < albite < K-feldspars.[3] Weathering of plagioclase minerals increases the supply of Ca in soils in a manner similar to that of weathering of K-feldspars in the clay, and silt fractions provides an important source of K in soils.

Micas are formed in soils from the parent material, and as they tend to weather to other minerals they are expected to prevail in younger, less weathered soils.[8] In young soils they tend to occur as discrete sand and silt particles, while in more weathered soils they are usually found in the clay fraction. Biotite is uncommon in most soils because it is transformed very easily to other secondary minerals, even under mild weathering conditions. Muscovite is commonly found in soils of advanced weathering. Like the other primary minerals in soils, mica becomes unstable in the weathering conditions at the soil surface. This is demonstrated by the fact that biotite is almost absent from the surface horizons of most, except very young, soils. Micas are considered as the precursors of several secondary 2 : 1 minerals and clay minerals in soil like illites, vermiculites, and smectites.[8,9]

Micas

Quartz

Micas are 2 : 1 phyllosilicate minerals formed by two silicon tetrahedral sheets and an aluminum octahedral sheet in between. One out of four Si in the tetrahedral sheet is replaced by Al and the high negative layer charge developed is balanced by nonhydrated interlayer cations that hold the layers tightly together.[7] The most abundant and important micas in soils (muscovite and biotite) have K as the interlayer cation, which upon weathering is released in the soil solution and serves as a plant nutrient.

Quartz is one of the seven polymorphs of SiO2, found in soils and sediments: quartz, cristobalite, tridymite, coesite, stishovite, lechatelerite (silica glass), and opal.[10] Quartz and disordered cristobalite are the typical forms found in soils while tridymite is present mostly in soils developed from siliceous volcanic rocks. Quartz belongs to the tectosilicate class of the silicate minerals (Table 1), which have each O of silicon tetrahedra linked to Si atoms of adjacent tetrahedra, forming a three-dimensional framework structure.


1086

Owing to its resistance to weathering, quartz is generally concentrated in the sand and silt fractions, and generally elluvial (surface) horizons are enriched owing to weathering and removal of less resistant minerals. Quartz is considered as inert sceletic material and exhibits very small to nil anion and cation exchange properties. In certain horizons, however, poorly crystalline quartz and other silica minerals may serve as cementing agents. In normal soils, quartz particles are also considered as important components of the soil structure because they are linked together or to clay domains through organic matter bridges.[11] Quartz is also a suitable mineral to assess parent material uniformity and degree of weathering because of its abundance, resistance to weathering, and immobility.

Minerals: Primary

PHOSPHATES Phosphate minerals comprise only a small part of the inorganic fraction of soils and their basic structural unit is the orthophosphate ion (PO43), which forms a tetrahedron where P is surrounded by four O2–. The high affinity of PO43– for cations, forms structures that are classified as framework, chain, and layer phosphates, similar to silicates.[13] Apatite [Ca10(PO4)6 (Cl,F,OH)2] is the most frequently reported phosphate of igneous origin and has been identified with photographic microscopy as discrete grains in the sand and silt fraction of a number of soils. Along with their weathering products, apatites are considered as a natural source of P for plants. CONCLUSIONS

OXIDES Iron Of the several iron oxides and hydroxides found in soils, magnetite (Fe3O4), is the only one of primary origin. The crystal structure of magnetite, like all iron oxides, consists of Fe ions surrounded by six O in a sixfold coordination. Magnetite occurs as black grains in the heavy fraction of the sand and silt fraction of soils and is of lithogenic origin.[12] As magnetite is found in the coarser soil fractions and is easily transformed to maghemite in the finer fractions, its influence on soil chemistry is minor. Titanium Rutile (TiO2) and ilmenite (FeTiO3) are the main primary titanium minerals found in soils. Rutile is tetragonal and consists of TiO6 polyhedra that share their opposite edges and form chains along the c-axis. Titanium oxides are common accessory minerals of igneous and metamorphic rocks and have been identified in the sand and silt fractions of soils originating from such rocks. Rutile develops a negative charge through adsorption of OH on surface Ti ions. This charge is pH dependent. Charge development plays an important role in sandy acid soils and influences physical and chemical properties such as formation of stable aggregates and retention of nutrients such as phosphate. Titanium oxides, in the similar way as zircon, are most frequently related to soil genesis studies. Their occurrence in the nonclay fraction in the different horizons of soils and the parent material provides information about the profile uniformity and soil development.


Primary minerals identified in soils belong mainly to the classes of silicates, oxides of Fe, Zr, and Ti, and phosphates and are usually found in the sand and silt fractions. In soils they are inherited from the parent material and their presence is predicated by the nature of the mineral and the intensity of weathering which results in the formation of secondary clay minerals in soils and the release of elements, essential for plant growth, in the soil solution. Therefore, primary minerals play a significant role in soil formation and in the long run they serve as sources of plant nutrients. ARTICLES OF FURTHER INTEREST Clay Minerals: Weathering and Alteration of, p. 281. Magnesium, p. 1045. Potassium, p. 1354. REFERENCES 1. Soil science society of America. Glossary of Soil Science Terms; Bigham, J.M., Ed.; SSSA: Madison, WI, 1997; 134 pp. 2. Shultze, D.E. An introduction to soil mineralogy. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 1–34. 3. Deer, W.A.; Howie, R.A.; Zussman, J. An Introduction to the Rock Forming Minerals; Longman Ltd.: Essex, England, 1992; 696 pp. 4. Huang, P.M. Feldspars, olivines and amphiboles. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 975–1050. 5. Churchman, G.J. The alteration and formation of soil minerals by weathering. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, FL, 2000; F3–F76.

Minerals: Primary

6. Allen, B.L.; Hajek, B.F. Mineral occurrence in soil environments. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 199–278. 7. Fanning, D.S.; Keramidas, V.Z.; El-Desoky, M.A. Micas. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 551–634. 8. von Reichenbach, H.G.; Rich, C.I. Fine grained micas in soils. In Soil Components. Part 2. Inorganic Components; Gieseking, J.E., Ed.; Springer: Berlin, 1975; 59–95. 9. Borchardt, G. Smectites. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 675–727.


1087

10. Drees, L.R.; Wilding, L.P.; Smeck, N.E.; Senkeyi, Abu L. Silica in soils: quartz and disordered silica polymorphs. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 913–974. 11. Emerson, W.W. The structure of soil crumbs. J. Soil Sci. 1959, 10, 235–244. 12. Schwertman, U.; Taylor, R.M. Iron oxides. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 379–438. 13. Lindsay, W.L.; Vlek, L.G.; Chien, S.H. Phosphate minerals. In Minerals in Soil Environments, 2nd Ed.; Series 1; Dixon, J.B., Weed, S.B., Eds.; SSSA: Madison, WI, 1989; 1089–1130.

Minerals: Secondary Deborah A. Soukup Geomatrix Consultants, Bakersfield, California, U.S.A.

INTRODUCTION There are three possible origins for minerals in soil environments including inheritance from parent materials, weathering transformations of existing minerals, and neoformation or crystallization from solution.[1] Secondary minerals are defined as minerals formed later than the rock enclosing them, usually at the expense of an earlier-formed primary mineral, as a result of weathering, metamorphism, or solution.[2] Alternatively, secondary minerals may be defined as recrystallized products of the chemical breakdown and=or alteration of primary minerals.[3] Secondary minerals are generally characterized by smaller particle size, because the particle size of primary minerals is decreased during weathering and release of soluble materials. Therefore, secondary minerals are typically principal components of the silt and clay fraction of soils.[4]

TYPES OF SECONDARY MINERALS IN SOIL The major secondary minerals and the soil orders in which they commonly occur are listed in Table 1. Primary minerals including feldspars, pyroxenes, amphiboles, micas, and primary chlorite may be altered to secondary minerals such as illite, vermiculite, clay chlorite, smectites, kaolinite, halloysite, and oxides of Fe and Al by the removal of silica and bases, and the addition of water. Additional details and the significance of these secondary minerals in soil environments are discussed in the following sections. Selected physical and chemical characteristics of the secondary minerals are summarized in Table 2. Iron and Aluminum Oxides The most common oxides of Fe and Al, typically reported in soils, are goethite, hematite, and gibbsite.[4,6,7] These minerals, sometimes referred to as sesquioxides, generally occur in soils subject to intense weathering, e.g., Oxisols and Ultisols. The sesquioxide minerals are amphoteric in character and exhibit variable charges. These minerals are also characterized by high surface charge density, high specific surface area, and high cation and anion adsorption capacity.[4–9] Additionally, sesquioxide minerals are reported to enhance soil aggregation. 1088 Copyright © 2006 by Taylor & Francis

Finely divided iron oxides are believed to act as binding agents among other soil particles and may result in cementation of these particles into large units.[7] Even at low concentrations, the iron oxides goethite and hematite play an important role in influencing soil color because of their pigmenting power.[7] Goethite minerals are naturally yellow, whereas hematite is bright red. In fact, the greater pigmenting power of hematite can mask the yellow color of higher concentrations of goethite. Gibbsite, in contrast, is colorless; however, its content is used in the oxidic ratio to indicate the relative degree of weathering: Oxidic ratio ¼

ð% extractable Fe2 O3 þ % gibbsiteÞ % clay

If the oxidic ratio is 0.2, the soil is considered to be highly weathered. Kaolinite and Halloysite Kaolinite is one of the most common clay minerals in soils, particularly those of warm, moist climates. Kaolinite is a 1 : 1 aluminosilicate mineral composed of one octahedral sheet stacked above one tetrahedral sheet. The two crystal units comprising one kaolinite particle are held together by hydrogen bonds, and the space between the structural layers, therefore, has a fixed dimension.[4,10] Halloysite is also a 1 : 1 aluminosilicate mineral with the same composition as kaolinite, except that halloysite may be hydrated and contain water between the structural layers. Both kaolinite and halloysite are products of acid weathering, but halloysite is formed more rapidly in soils of volcanic origin. Kaolinite and halloysite typically have low surface areas and low cation- and anion-exchange capacities. Isomorphous substitution within the crystal is limited, contributing to the low permanent charge. However, kaolinite and halloysite may develop variable or pH dependent negative charge because of the dissociation of protons for exposed OH groups.[4,10] The typical range in cation exchange capacity (CEC) for kaolinite and halloysite is from 1 to 10 cmol(þ)=kg. Kaolinite-containing clays are used extensively in the production of brick, sewer pipes, and drain tiles. Kaolinite is also used in the ceramic industry, because of its low expansion and contraction capacity. Additionally, Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042714 Copyright # 2006 by Taylor & Francis. All rights reserved.

Minerals: Secondary

1089

Table 1 Ideal chemical formula of secondary minerals commonly occurring in soils Mineral name Oxides of Fe and Al Goethite Hematite Gibbsite Aluminosilicate clays Kaolinite Halloysite Illite Chloritea Vermiculite Smectite Hydroxy-interlayer Vermiculite and Smectite Carbonate and sulfate minerals Dolomite Calcite Gypsum

Chemical formula

Major occurrences

FeOOH Fe2O3 Al(OH)3

Oxisols, Ultisols Oxisols, Ultisols Oxisols, Ultisols

Al2Si2O5(OH)4 Al2Si2O5(OH)4 2H2O K(Si3Al)(AlMg)2O10(OH)2 [(M2þ,M3þ)3(SiAl)4(Al2)O10 (OH)2]x[(M2þ,M3þ)3(OH)6]xþ

Ultisols, Oxisols, Alfisols Andisols Alfisols, Mollisols Common accessory mineral in many soils

Mg(SiAl)4(Al2)O10(OH)2 Na(Si4)(AlMg)3O10(OH)2 Chemical composition highly variable

Common in many soils Vertisols, Mollisols, Alfisols Wide geographic distribution but most abundant in Ultisols and Alfisols

CaMg(CO3)2 CaCO3 CaSO42H2O

Aridisols, Mollisols, Alfisols, arid Entisols Aridisols, Mollisols, Alfisols, arid Entisols Aridisols

M2þ and M3þ represent cations with charges of þ2 and þ3, respectively.

a

kaolinite is used for the production of false teeth, paper making, and in the pharmaceutical industry.[4,10] Illites Illites are micaceous clays and have been referred to as hydrous micas and hydrobiotite.[4] However, illite differs from micas, such as muscovite, in that it

contains more SiO2 and less K. Illites are 2 : 1 aluminosilicate minerals consisting of one octahedral sheet sandwiched between two tetrahedral sheets. The 2 : 1 layers are bound together strongly by K ions, and thus, illites are nonexpansive and do not exhibit shrink–swell behavior upon wetting and drying.[4,11] Illite typically weathers to smectite under conditions of high precipitation, because of the loss of interlayer K.

Table 2 Selected physical and chemical characteristics of secondary minerals commonly occurring in soils Mineral name Oxides of Fe and Al Goethitea Hematitea Gibbsiteb Aluminosilicate clays Kaolinite Halloysite Illite Chlorite Vermiculite Smectite Hydroxy-interlayer Vermiculite and Smectite Carbonate and Sulfate minerals Dolomitec Calcitec Gypsumc a

See Refs.[5,7,8] for additional information. See Refs.[5,6,9] for additional information. c See Ref.[16] for additional information. b


Surface area (m2/g)

Cation exchange capacity (cmolc/kg)

50–200 50–200 58>600

3 3 1–3

5–39 21–43 80–150 25–150 600–800 600–800 25–150

2–15 10–60 20–40 10–40 100–200 80–150 CEC varies depending on degree of filling of interlayer space

1090

Vermiculites Vermiculite is also a 2 : 1 aluminosilicate mineral with Mg occupying the octahedral positions between the two tetrahedral sheets.[4,12] Vermiculites are common weathering products of micas in well-drained soils and are characterized by high CECs ranging from 144 to 182 cmol(þ)=kg.[12] These minerals exhibit wedge zones, attributed to marginal curling of layers on the mineral surface.[4] These wedge zones provide partially enlarged interlayer spacings in which organic matter may be entrapped, or for fixation of K, NH4þ, and other cations. The high fixation capacity of many soils for K and NH4þ is primarily attributable to the presence of vermiculite. Vermiculite is also widely used as a potting medium in nurseries and in cat litter.

Smectites Smectites are a group of expansive, 2 : 1 aluminosilicate minerals, commonly occurring in soils with impeded drainage and=or in soil environments characterized by high Si and basic cation activities.[13] Minerals within the smectite group are classified based upon the location of the layer charge, either within the tetrahedral or octahedral sheets. In montmorillonite, substitution of Mg for Al in the octahedral sheet produces the layer charge. The layer charges of beidellite and nontronite, two other important smectite minerals, are concentrated in the tetrahedral sheets. The predominant octahedral cations in beidellite and nontronite are Al and Fe, respectively.[13] Smectites are generally responsible for the high shrink– swell conditions in soil and result in more damage annually than any other natural disaster including earthquakes and floods. Annual estimates of damage to residential structures in the U.S. were approximately 798.1 million in 1970 and are projected to increase to 997 million by the year 2000.[14] These figures increase by a factor of 2–3 if damage to industrial and commercial buildings and transportation infrastructure are included.

Hydroxy-Interlayered Vermiculite and Smectite Hydroxy-interlayered vermiculite (HIV) and smectite (HIS) are believed to occur in soils as weathering products derived from chlorite or micas or from the deposition of hydroxy-Al or -Fe polymeric components within the interlayer spaces of vermiculite and smectite.[15] HIV and HIS are widely distributed geographically and are found in several soil orders, although they are most commonly reported in Ultisols and Alfisols.[15] There is significant variability in the composition of HIV and HIS, because the composition is dependent


Minerals: Secondary

on the basic 2 : 1 mineral structure and the type and amount of hydroxy-interlayer material within the interlayer spaces. The presence of hydroxy-Al and -Fe interlayer components may have a profound effect on some of the physical and chemical properties of HIV and HIS, including swelling, CEC, and adsorption of metals and anions. Many investigators have reported that formation of hydroxy-Al and -Fe interlayers resulted in reduced swelling and dispersion of smectites. Hydroxy-Al polymers were more effective than hydroxyFe polymers in reducing swelling, presumably because the Al interlayer components are more uniformly distributed within the interlayer space, and their greater relative stability.[15] The CEC of HIV and HIS may also be reduced as the interlayer space is filled. The primary mechanisms proposed for the observed charge reduction in HIV and HIS compared to the noninterlayered end-members are precipitation of Al on surfaces and=or into interlayer spaces, sterically blocking exchange sites, and adsorption of positively charged hydroxy-Al polymers that are nonexchangeable.[15] Hydroxy-Al interlayers can also prevent fixation of Kþ, NH4þ, Csþ, and Rbþ because the interlayer material inhibits collapse of the layers about these ions. Finally, the presence of hydroxy interlayers in 2 : 1 clay minerals significantly enhances the adsorption of anions, such as P.[15] The increase in anion adsorption in hydroxy-interlayered clays occurs because anions are only adsorbed at edge sites in minerals not containing hydroxy-interlayers. Chlorites Chlorites are hydrated Mg and Al silicates that are commonly green in color and resemble micas in appearance.[4,15] Structurally, chlorite is a 2 : 2 aluminosilicate mineral consisting of one octahedral sheet and two tetrahedral sheets with the interlayer space occupied by Mg(OH)2 or brucite layers. The replacement of Mg by Al in the brucite layers creates enough positive charge to nearly neutralize the negative charges. Thus, chlorite has little or no charge and a small CEC. Chlorites are less stable than most of the other clays in acidic environments and are subject to rapid weathering.[4,15] Carbonate and Sulfate Minerals Calcium and magnesium carbonate minerals, including calcite and dolomite, may originate from several sources or combinations of sources, either directly in the form of carbonates, or by a solution-precipitation mechanism.[16] Dissolution of more soluble Ca-bearing minerals, such as gypsum, may result in the subsequent precipitation of calcite. Carbonates may also be formed by the reaction of Ca ions in rainwater and surface water with

Minerals: Secondary

CO2-charged H2O in the soil. Coatings of calcium carbonate may eventually accumulate and result in cementation in some soils. Movement of clay in calcareous soils has also been reported to be restricted because of flocculation of the silicate clays by Ca2þ.[16] Gypsum may be precipitated in soil from surface or subsurface waters rich in Ca and SO4. It may also be formed through the oxidation of pyrite and subsequent reaction of acid sulfate with CaCO3.[16] Gypsum has frequently been used to reclaim sodic soils because of its high solubility. However, numerous cases of soil subsidence and deterioration of concrete due to the presence of gypsum have been reported.

CONCLUSIONS In summary, it should be noted that the information presented in this entry on the formation and characteristics of common secondary minerals in soils is a broad overview that simplifies many relevant details. The actual alteration pathways occurring in the soil environment are very complex and are influenced by a variety of abiotic and biotic factors. The secondary minerals discussed in this entry may also occur in soils as a result of inheritance from parent materials. Additionally, it may be extremely difficult in some soil environments to determine the precise origin of a given mineral.

REFERENCES 1. Allen, B.L.; Hajek, B.F. Mineral occurrence in soil environments. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 199–278. 2. Bates, R.L.; Jackson, J.A. Dictionary of Geological Terms 3rd Ed.; Doubleday: New York, 1984. 3. Brady, N.C.; Weil, R.R. The Nature and Properties of Soils, 12th Ed.; Prentice-Hall: Upper Saddle River, 1999. 4. Tan, K.H. Environmental Soil Science, 2nd Ed.; Marcel Dekker: New York, 2000; 28–79.


1091

5. McBride, M.B. Environmental Chemistry of Soils; Oxford University Press: New York, 1994. 6. Hsu, P.H. Aluminum hydroxides and oxyhydroxides. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 331–378. 7. Schwertmann, U.; Taylor, R.M. Iron oxides. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 379–438. 8. Bigham, J.M.; Fitzpatrick, R.W.; Schulze, D.G. Iron oxides. In Soil Mineralogy with Environmental Applications, Dixon, J.B., Schulze, D.G., Eds.; Soil Science Society of America: Madison, 2002; 323–366. 9. Huang, P.M.; Wang, M.K.; Ka¨mpf, N.; Schulze, D.G. Aluminum hydroxides. In Soil Mineralogy with Environmental Applications; Dixon, J.B., Schulze, D.G., Eds.; Soil Science Society of America: Madison, 2002; 261–289. 10. Dixon, J.B. Kaolin and serpentine group minerals. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 467–525. 11. Fanning, D.S.; Keramidas, V.Z.; El-Desoky, M.A. Micas. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 551–634. 12. Douglas, L.A. Vermiculites. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 635–674. 13. Borchardt, G. Smectites. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 675–727. 14. Reid, D.A.; Ulery, A.L. Soil Mineralogy with Environmental Applications; Dixon, J.B., Schulze, D.G., Eds.; Soil Science Society of America: Madison, 2002; 467–499. 15. Barnhisel, R.I.; Bertsch, P.M. Chlorites and hydroxyinterlayered vermiculite and smectite. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 729–788. 16. Doner, H.E.; Lynn, W.C. Carbonate, halide, sulfate, and sulfide minerals. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America: Madison, 1989; 279–330.

Minerals: Solubility Wayne P. Robarge North Carolina State University, Raleigh, North Carolina, U.S.A.

INTRODUCTION Minerals are found in almost all soils and the precipitation and dissolution reactions of minerals are an integral component of the biogeochemical cycles in soils. A mineral is considered a naturally occurring substance, inorganic in composition, with a definite chemical composition and an ordered atomic arrangement.[1,2] Minerals in soils are divided into two broad categories: primary and secondary minerals.

MINERALS Primary Minerals A primary mineral is considered one that has not been altered chemically since its deposition and crystallization as a result of large-scale geological processes.[1] Soils inherit primary minerals from the parent material from which the soil is derived.[3] Chemical weathering of primary minerals (a dissolution reaction) represents the release of plant nutrients such as Mg, Ca, K, P, Fe, Mn, and B for plant uptake and recycling in soil–plant ecosystems. Primary minerals are also a source of trace metals, which can either be essential or toxic to life. As most primary minerals are formed by tectonic activities, they are inherently unstable in soils and will dissolve.

Additional secondary minerals found in soils include oxides, sulfates, sulfides, carbonates,[5] and phosphates.[6] In arid environments, the presence of sulfates and carbonates is common, but they are generally absent from well-leached soils in humid environments. Most oxides are stable in well-aerated soils and are considered the end product of the decomposition of primary and secondary minerals in highly weathering environments. However, the presence of organic acids derived from the decomposition of plant material can dissolve most oxides in acid soils. Certain metal oxides based on Fe and Mn are readily soluble under reducing conditions induced in poorly drained soils or when soils are temporarily saturated with water. Sulfides are generally not stable in well-aerated soils, due to the conversion of the sulfide anion to sulfate in the presence of oxygen. Sulfides are common in anaerobic soils where there is a source of sulfur either from plant residues or sulfate secondary minerals. Phosphates encompass a wide range of minerals with differing solubility. In neutral and basic soils, the Ca-based phosphates, known as apatites, are relatively insoluble and can remain in soils for long periods of time. In acid soils, Al- and Fe-based phosphates are considered stable, especially in agricultural soils receiving P-containing fertilizers. The names and chemical formula of some common primary and secondary minerals are given in Table 1.

Secondary Minerals

DISSOLUTION REACTIONS OF MINERALS

Secondary minerals form from the decomposition and restructuring of primary minerals or from precipitation reactions (formation of a new mineral) involving chemical constituents of primary minerals released during dissolution.[1] Most secondary minerals in soils are phyllosilicates: silicate-based minerals having sheets of silicate anions linked to sheets of cations (usually Al) to form a layer.[2] The cations in phyllosilicates also have reacted with oxygen such that the sheets of silicate anions and cations are linked together by sharing oxygen atoms. This yields a very stable structure that resists weathering. Phyllosilicates can be formed in soils, or are inherited in the soil parent material, having been formed elsewhere as part of the weathering of primary minerals.[4]

Solubility of Minerals


The stability of primary and secondary minerals in soil, and therefore their solubility, is directly dependent on their chemical structure and the nature of the weathering environment.[3] Differences in chemical bonding among minerals influences their shape, often exposing surfaces or defects in the minerals that react more readily with water or serve as points of attack for other constituents in the soil solution. The nature of the weathering environment is expressed through the composition of the soil solution [pH, ionic strength (presence of other cations and anions), presence of ligands, reductants, oxidants, or ions that can adsorb onto the mineral surface] and soil drainage, which Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006571 Copyright # 2006 by Taylor & Francis. All rights reserved.

Minerals: Solubility

1093

Table 1 Names, chemical formula, and dissolution reaction for common minerals Name

Chemical formula

Dissolution reaction

Carbonates Calcite Dolomite Siderite

CaCO3 MgCa(CO3)2 FeCO3

Ca2þ þ CO32 Mg2þ þ Ca2þ þ 2CO32 Fe2þ þ CO32

Sulfates Gypsum Alunite

CaSO4 2H2O KAl3(SO4)2(OH)6

Ca2þ þ 2SO42 þ 2H2O Kþ þ 3Al3þ þ 2SO42 þ 6OH

Phosphates Variscite Strengite Hydroxyapatite Fluorapatite

AlPO4 2H2O FePO4 2H2O Ca10(OH)2(PO4)6 Ca10F2(PO4)6

Al3þ þ Fe3þ þ 10Ca2þ 10Ca2þ

Sulfides Alabanite Pyrite Black-metacinnibar

MnS FeS2 b-HgS

Mn2þ þ S2 Fe2þ þ S22 Hg2þ þ S2

Oxides/Hydrousoxides Quartz Gibbsite Hematite Goethite

a-SiO2a g-Al(OH)3 a-Fe2O3a a-FeOOHa

H4SiO4 Al3þ þ 3OH Fe3þ þ 3OH Fe3þ þ 3OH

Silicates Olivine Diopside Pyroxene Albite Microcline Anorthite

Mg1.6Fe0.4SiO4 CaMg(SiO3)2 CaAl2SiO6 NaAl2Si3O8 KAlSi3O8 CaAl2Si2O8

1.6Mg2þ þ Fe2þ þ H4SiO4 Mg2þ þ Ca2þ þ 2H4SiO4 Ca2þ þ 2Al2þ þ H4SiO4 þ 2H2O Naþ þ Al2þ þ 3H4SiO4 Kþ þ Al3þ þ 3H4SiO4 Ca2þ þ 2Al3þ þ 2H4SiO4

Phyllosilicates Muscovite Kaolinite

KAl2(AlSi3O10)(OH)2 Al2Si2O5(OH)4

Kþ þ 3Al3þ þ 3H4SiO4 2Al3þ þ 2H4SiO4 þ H2O

a

PO43 þ 2H2O PO43 þ 2H2O þ 2OH þ 6PO43 þ 2F þ 6PO43

Dissolution of Si-containing minerals often requires H2O to form the silicate anion: H4SiO4.

influences the rate at which constituents released from the mineral surfaces are transported from the local environment.[7] Acid (Hþ) or base (OH) (and even water molecules) destroy mineral structures by migrating into the interior of the mineral and breaking the chemical bonds between the mineral’s anions and cations. Ligands are inorganic or organic molecules that have the ability to bind tightly with the cations in minerals and help to remove them from a mineral surface. Reductants and oxidants cause the cations in minerals to change their charge (reductants reduce charge, oxidants increase charge). When a cation within a mineral structure changes charge, it weakens the original structure allowing the attack by acids, bases, or water molecules to be more successful. Soil drainage is important in mineral dissolution because as the original ionic constituents of the mineral accumulate in the surrounding soil solution, mineral dissolution will eventually stop or slow significantly. Thus primary and secondary minerals are often found


in weathering environments that would favor their relatively rapid dissolution because of poor drainage in the soil. Chemical theories derived to predict the behavior of minerals in soils and that attempt to account both for differences in chemical structure and the nature of the soil weathering environment have evolved from consideration of either macroscale or microscale processes.[8] Theories that describe macroscale processes are based on the science of thermodynamics and allow prediction of mineral solubility.[9] The power of this approach is that a scientist can attempt to predict the presence of a mineral without actually seeing it in the soil. The disadvantage of this approach is that it cannot easily predict the rate at which minerals will dissolve. Theories that focus on microscale processes assume specific chemical reaction mechanisms occur at mineral surfaces and are more successful in predicting the rate of mineral dissolution due to changes in the composition of the surrounding soil solution.[10]

1094


Unfortunately, microscale processes require information that can usually only be generated under laboratory conditions, which, at best, only approximate the actual weathering environment in the soil. A knowledge of both macroscale and microscale processes is important in understanding the solubility of minerals.[8] Macroscale Processes Theories to describe mineral dissolution based on macroscale processes assume that the minerals, which may be present in a soil, are only sparingly soluble in the presence of water. In other words, only a tiny fraction of the mineral dissolves at any one time. The amount that dissolves is in proportion to the chemical formula for the mineral. Chemical analysis of the soil solution in contact with the soil minerals, therefore, should yield information about the type of minerals present. For example, the dissolution reaction for the secondary mineral gibbsite can be written as: AlðOHÞ3ðsÞ þ 3HðaqÞþ () AlðaqÞ3þ þ 3H2 OðlÞ

ð1Þ

where the subscript (s) refers to the solid mineral, (aq) to the ion in solution, and (l) for water. The () symbol reflects the assumption that the solid gibbsite mineral is controlling the concentration of Al3þ ions in the soil solution. Thermodynamics predicts that if gibbsite is acting to control the composition of the soil solution, then the product of the concentrations of the dissolved ions in the soil solution is a constant (often referred to as the solubility product; Kso). The Kso of a mineral is in theory a unique quantity that is specific to the chemical composition and chemical structure of the mineral.[9] A visual representation of the ability to predict the presence of a mineral based on the concept of the Kso is shown in (Fig. 1) for the mineral gibbsite, where the product of the concentrations of the dissolved ions in soil solution represented by Eq. (1) has been written in the form of a straight line: logðAl3þ Þ ¼ K0 3pH

mineral (a precipitation reaction). If the data points fall in a region below the lines (region of undersaturation), conditions in the soil are considered too severe for the mineral to survive for long periods of time and unlikely that the mineral would be present in the soil.[12] The four straight lines illustrate that solubility is not just dependent on the chemical composition of a mineral, but also on its degree of crystallinity (amorphous is more soluble than synthetic) and size (microcrystalline is more soluble than more normal sized particles).[13] The concept of the solubility product for minerals and the use of diagrams (often termed activity-ratio or activity–activity diagrams) as illustrated in Fig. 1 provides a useful means of predicting mineral behavior in soils, often by delineating boundaries for possible reactions between several minerals.[13–15] The approach is dependent, however, on the assumption that the rate at which minerals dissolve is not limiting and thereby influencing the concentrations of ions in solution. Predicting rates of dissolution in specific soil environments is best viewed as a microscale process.

ð2Þ

The term log(Al3þ) represents the activity of Al3þ in solution (a measure of concentration), K0 is a constant whose value is dependent on Kso, pH is the activity of Hþ in solution, and the number 3 represents the fact that it takes 3 Hþ ions to release one Al3þ ion. Experimental data points plotted on such graphs, which fall along the linear lines, are taken to indicate the presence of the mineral in the soil. If the data points fall in a region above the lines (region of supersaturation), conditions are favorable for the formation of the


Fig. 1 Activity-ratio diagram for synthetic gibbsite, natural gibbsite, microcrystalline gibbsite, and amorphous aluminum hydroxide. Kdiss values taken from Ref.[11].

Microscale Processes It is now generally accepted that the rate of dissolution of minerals in soils is controlled by chemical reactions along mineral surfaces at distinct surface features that can include dislocations, etch pits, and micropores.[16] The presence of surface features like dislocations, etch pits, and micropores is taken as an indication of a population of surface reactive sites where dissolution reactions are favored above those possible along the


total surface area of a mineral.[8] The dissolution reaction at the surface reactive sites is considered a two-step process. The first step involves the relatively rapid formation of a surface species as the result of the reaction with entities in soil solution such as Hþ, OH, H2O molecules, and inorganic and organic ligands. The second step, and the one that is considered rate-limiting, involves the transformation of the surface species into an activated complex with the eventual release of constituent ions to the bulk solution.[7] Dissolution at the microscale, therefore, can be considered as a function of the reactive surface area of the mineral, the nature of the inherent chemical bonding structure that comprises the mineral, and the presence of various chemical species in the soil solution. Thus the dissolution rate of a mineral is not a fixed numerical quantity but dependent upon a number of variables which may change with time due to changes in the physical size of the mineral itself, as well as changes in the surrounding soil solution.[8] It is possible to calculate approximate rates of dissolution for various classes of primary and secondary minerals, but local conditions within the soil should also be considered when predicting mineral solubility.

REFERENCES 1. S374 glossary of soil science terms committee. In Glossary of Soil Science Terms; Soil Science Society of America: Madison, WI, 1997; 100 pp. 2. Bragg, L.; Claringbull, G.F.; Taylor, W.H. Crystall Structures of Minerals: The Crystalline State; Cornell University Press: Ithaca, New York, 1965; Vol. IV, 409 pp. 3. Allen, B.L.; Hajek, B.F. Mineral occurrence in soil environments. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America Book Series Number 1; Soil Science Society of America: Madison, Wisconsin, 1989; 199–278. 4. White, A.F. Chemical weathering rates of silicate minerals in soils. In Chemical Weathering Rates of Silicate Minerals; White, A.F., Brantley, S.L., Eds.; Reviews in Mineralogy; The Mineralogical Society of America: Washington, DC, 1995; Vol. 31, 407–461.


1095

5. Doner, H.E.; Lynn, W.C. Carbonate, halide, sulfate, and sulfide minerals. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America Book Series Number 1; Soil Science Society of America: Madison, Wisconsin, 1989; 279–330. 6. Lindsay, W.L.; Vlek, P.L.G.; Chien, S.H. Phosphate minerals. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Science Society of America Book Series; Soil Science Society of America: Madison, Wisconsin, 1989; 1089–1130. 7. Robarge, W.P. Precipitation/dissolution reactions in soils. In Soil Physical Chemistry, 2nd Ed.; Sparks, D.L., Ed.; CRC Press: Boca Raton, 1999; 193–238. 8. Lasaga, A.C. Fundamental approaches in describing mineral dissolution and percipitation reactions. In Chemical Weathering Rates of Silicate Minerals; White, A.F., Brantley, S.L., Eds.; Reviews in Mineralogy; The Mineralogical Society of America: Washington, DC, 1995; Vol. 31, 23–86. 9. Sposito, G. The Thermodynamics of Soil Solutions; Oxford University Press: London, 1981; 223 pp. 10. Lasaga, A.C. Kinetic Theory in the Earth Sciences; Princeton University Press: Princeton, New Jersey, 1998; 811 pp. 11. Johnson, N.M.; Driscoll, C.T.; Eaton, J.S.; Likens, G.E.; McDowell, W.H. ‘Acid rain,’ Dissolved aluminum and chemical weathering at the hubbard brook experimental forest, New Hampshire. Geochim. Cosmochim. Acta 1981, 45 (9), 1421–1437. 12. Lindsay, W.L. Chemical equilibria. In Soils; John Wiley & Sons: New York, 1979; 449 pp. 13. Dove, P.M. Kinetic and thermodynamic controls on silica reactivity in weathering environments. In Chemical Weathering Rates of Silicate Minerals; White, A.F., Brantley, S.L., Eds.; Reviews in Mineralogy; The Mineralogical Society of America: Washington, DC, 1995; Vol. 31, 235–290. 14. Garrels, R.M.; Christ, C.L. Solutions, Minerals and Equilibria; Harper and Row: New York, 1965; 450 pp. 15. Stumm, W.; Morgan, J.J. Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters, 3rd Ed.; John Wiley & Sons, Inc.: New York, 1996; 1022 pp. 16. White, A.F.; Brantley, S.L. Chemical weathering rates of silicate minerals: an overview. In Chemical Weathering Rates of Silicate Minerals; White, A.F., Brantley, S.L., Eds.; Reviews in Mineralogy; The Mineralogical Society of America: Washington, DC, 1995; Vol. 31, 1–22.

Modern Civilization and Soils Ken R. Olson University of Illinois, Urbana, Illinois, U.S.A.

INTRODUCTION In many parts of the world, contemporary use of soils is not greatly different from ancient use of soils.[1] The archeological record shows that some ancient people abused their soils and that their civilizations were disrupted by the ecological and environmental consequences. Some ancient populations, such as the Middle Mississippian culture near Cahokia, Illinois, were larger in population in 1200 A.D. than their modern descendents.[2] The archeological record gives us reasons to ponder, whether contemporary societies are using their resources wisely.[1] People should seek explanations for ancient population declines and shifts so they might avoid similar fates, because an important factor in the continuing prosperity of a nation is the nature of the husbandry of the soil and land resources through centuries of use and occupation. The Middle Mississippian Indians settled in an area near Cahokia, in southwestern Illinois (adjacent to the current city of St. Louis, Missouri) in approximately 500 A.D.[3–5] By 1350 A.D. the site was abandoned after peaking in the Lohmann phase (1050–1100 A.D.) and the Stirling phase (1100–1200 A.D.).[2,5–7] During the later years of settlement, the population grew to 20,000 people, making it the largest city in the U.S.A. until the European settlement of Philadelphia and Pennsylvania exceeded 20,000 people in 1830. The Middle Mississippian people settled on the stream and lacustrine deposits in the American Bottoms of the Mississippi River valley. Previously, it would not have been possible to feed such a large population by hunting and fishing alone. The Middle Mississippi Indians supplemented their diet by cultivating maize (corn) and garden crops on the silty soils (Fluvaquents) on floodplains.[8] They could then store sufficient food for the winter season to feed the entire population. The people also cut down the adjacent forests for firewood and making structures and fences. This deforestation caused accelerated erosion of the steep, silty soils (Hapludalfs) on nearby hillsides. As a result, wood had to be transported to greater and greater distances over time. These Middle Mississippian Indians were also earthen mound builders. They created thousands of earthen burial and ceremonial (platform) mounds. The earthen mounds were built on clayey soils (Fluvaquents). They dug out the silty, clayey, and sandy Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042715 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

stream deposits from the Mississippi River valley and carried the soil materials on their backs in sacks to the top of the mounds.[2] Later, the original borrow areas were filled to build a large plaza by digging and transporting more soil materials in sacks from borrow areas farther away. Some of these mounds (almost 800 years old) still exist. The largest mound is called Monks Mound, which is still 30 m high with a rectangular base of approximately 320 m 294 m. War, fire, food shortages, and disease were the most likely causes for the collapse of the Middle Mississippian civilization, but the final causes for the collapse are still the subject of considerable debate. It is difficult to understand how the wastes from 20,000 people could be disposed off without polluting the adjacent water and land. Waste disposal problems could have contributed to the contamination of the water supplies, especially the water in the borrow areas. Over time, the demand for maize without crop rotation and nutrient additions could have resulted in soil depletion and reduced crop yields, which could have contributed to food shortages. Deforestation of forests, overgrazing of pastures and ranges, and overplowing of fields may have accelerated erosion in many ancient civilizations. Accelerated soil erosion may have diminished the ability of many farmers worldwide to produce satisfactory yields and maintain a reasonable standard of living.[9] Tillage utilized in agriculture often accelerates soil loss from water and wind erosion.[10] For survival and well-being, governments throughout the world need to respond to the challenges of the stabilizing erodible soils.

MODERN CIVILIZATIONS World demand for food climbs higher each year as a result of both population growth and rising expectations. As a nation’s standard of living increases, diets often contain more of meat. This change may entail growing of feed crops, such as corn and soybeans, which lead to more erosion than of small grains and forage.[11] In the Third World, farmers are being pushed onto steeply sloping, erodible lands that are rapidly losing their topsoil. Even in the American Midwest, many farmers have abandoned more biologically 1099

1100

Modern Civilization and Soils

diverse and soil-conserving rotations that include forage and small grains (oats and wheat) in favor of continuous corn, or a corn and soybean rotation. In other parts of the world, farming has extended into semiarid regions where land is vulnerable to soil loss from wind erosion when plowed.[12] The loss of topsoil affects the ability of soil to grow food crops, by decreasing the availability of nutrients. This results in degrading the soil’s physical condition by increasing the need for fertilizers and irrigation, which, in turn, increases the cost of production. The loss of topsoil is a quiet crisis in the world economy, because unlike natural disasters such as earthquakes or volcanic eruptions, it is largely a human-made crisis that unfolds gradually.[11] Soil degradation is more severe in sub-Saharan Africa than other tropics.[12] This is where soil degradation has severely reduced the soil’s potential capacity to produce food, feed, and fiber. This is a result of physical, biological, and chemical deterioration due to erosion, desertification, laterization (formation of a stone-like layer high in iron), salt accumulation, excessive leaching and acidification, and nutrient imbalance. Prior to the European settlement, many northcentral U.S. soils were developed under prairie grass on nearly level plains and under forests on rolling hillslopes. European settlement and agricultural technology had significant impacts on north-central U.S. soils Table 1.

During the Dust Bowl of the 1930s, severe droughts in the Great Plains of the U.S.A. resulted in crop failures and major dust storms. In 1935, with dust settling in the congressional offices in Washington, D.C., the Congress passed a law, creating the Soil Conservation Service to address the erosion problem. At that time, soil erosion of row cropland became so severe, and farm debt became so high, that lands were abandoned in many states such as Illinois, Indiana, Kentucky, Tennessee, Oklahoma, and Missouri. Numerous research stations and demonstration farms were established in the northcentral U.S.A. to test and disseminate information on how to restore soils and improve crop yields and farm income on hilly, highly eroded soils. Management techniques that helped achieve these goals included the removal, crushing, and application of limestone from nearby hills to highly eroded soils with acidic subsoils. Rock phosphate was added to the phosphorus-poor soils and forage crops were seeded as pasture for beef and sheep, reducing the need for a yearly tillage. Gullies were filled on the steep hillslopes as well. In the 1950s, nitrogen fertilizer became available at low cost and the increasing acreage planted with soybeans and corn increased erosion. In the 1960s, the moldboard plow was partially replaced with a chisel plow, which reduced soil loss by leaving plant residue on the surface (conservation or mulch tillage). In the 1970s, no-till planting of limited

Table 1 Agricultural technology impacts of European settlement on north-central U.S.A. Timeline

Agricultural technology events related to European settlement

1814

The introduction of the cast-iron plow made it possible to plow the forest land but not the wet, sticky prairie soils. Consequently, settlement in the early 1800s along the Ohio, Mississippi, and Missouri Rivers resulted in the clearing of forest and cultivation of sloping land which was easily eroded

1837

The John Deere Company started mass producing a steel plow which could be used to plow the prairie lands. Shortly thereafter, some of the agriculture from the highly eroded forest soils shifted or moved to the prairie soils

1850s

The steam engine was introduced which helped with the wheat harvest and increased wheat production. In 1879 the U.S. Drainage Act and Levee Act allowed the drainage (initially surface ditches and later tile systems) and farming of swampy prairie areas in states like Illinois

1890–1950

Gasoline tractors replaced horses resulting in a decline in the acreages of oats and hay

1920s

Agronomic research resulted in better oil seed varieties of soybean. Crop rotations which had previously included corn planted in rows approximately 1 m wide, drilled wheat and oats in very narrow rows approximately 7.5 cm wide, and perennial hay was replaced by an intensive rotation of corn and soybeans. Over time, the soybeans were grown in wide rows for grain instead of being closely grown for forage; consequently, accelerated soil erosion occurred



acreages of corn and soybean further reduced tillage and soil erosion, but often with increased use of herbicides. In the 1980s, ridge tillage, which left residues between the rows at planting, to reduce erosion, became more popular. Row cultivation was used to control weeds between the rows and to maintain ridges. The U.S. Food Security Act of 1985 resulted in millions of hectares of highly erodible lands being put into forages or trees for 10 years or more, which further helped reduce soil movement from the row cropland. Many states developed conservation programs in the 1980s that required soil loss to be reduced over a specific time, to the tolerable soil loss level, ranging from 2 to 11 metric tons per hectare per year.

EFFECTS OF AGRICULTURE ON SOILS Sheet and rill erosion rates on cropland have been declining in U.S.A.[13] Under current production of agricultural systems, however, 23% of U.S. cropland is still eroding at rates higher than that of the tolerable soil loss.[14,15] The effects of erosion on cropland soils depend, in part, on the characteristics of the soils, and the landscapes in which they occur. Even under favorable climatic conditions, topography, and parent materials, soil formation is a slow process, requiring hundreds of years to develop topsoils and subsoils. Recently adopted conservation management improvements, advances in conservation tillage technology, and land-use changes have contributed to sheet and rill erosion-rate reductions; however, still there can be a net soil loss and reduced longterm soil productivity even under no-till systems. When considered in the context of long-term productivity, erosion rates alone are not the only indicator of soil degradation. Degradation of soil structure and aggregation or loss of irreplaceable soil attributes are much more serious in certain soils than others when compared at the same erosion rates. Documented on-farm and off-site damages include loss of soil, sedimentation in streams, reservoirs, and fields, loss of nutrients, soil property changes, soil productivity losses, and air, soil, and water pollution. Current government programs either provide incentives or require soil loss to be held within tolerable levels. On average, the total soil loss resulting from wind and water activity has been declining, in part, as a consequence of government programs. However, erosion and sediment deposition are locally and regionally significant and, thus, still perceived by the scientific and nonscientific communities as a serious threat to our economy and environment.


1101

EFFECTS OF MODERN URBANIZATION AND INDUSTRIALIZATION ON SOILS For the past 200 years, as European settlement increased in the U.S.A., the disturbance and alteration of the soils have accelerated. Many millions of hectares in the urban landscape have been altered or destroyed for constructing buildings, transportation, or mineral extractions. Vast areas of soils near urban centers have been excavated or disturbed for the construction of buildings, roads, railroads, cemeteries, canals, utility lines, and landfills. In addition, there has been surface and subsurface mining for various minerals, including coal, limestone, clay, sand, and gravel apart from waste disposal. Topsoils and even subsoils are often removed and sold from constructions sites. Topsoil has become a commodity that is of value and is often transported and sold to the highest bidder. In the urban landscapes, soils (often both topsoil and subsoil) have been removed, exposing the dense parent materials, which are often built on and then landscaped using a thin layer of sod. This sod has difficulty in growing roots into the dense parent material. Often, soils are no longer present and cannot accept or hold the rainwater on the landscape. Runoff is much greater than it was prior to the alteration. The lack of soil water storage and covering of other soils in the landscape by roofs, sidewalks, driveways, and roads has resulted in more runoff and increased flooding. As humankind have tried to clean up the air and water, the industrial and urban pollutants, urban and industrial wastes, and sewage sludges have either been consolidated or buried under the soil materials, or transported and spread on soils as amendments. Fortunately, environmental regulations no longer allow urban and industrial wastes and sewage sludges to be transferred to the air, rivers, lakes, or oceans.

CONCLUSIONS Contemporary use of soils is not greatly different from ancient use of soils. Many millions of hectares in the modern urban landscape have been altered or destroyed for construction of buildings, transportation, or mineral extractions. The archeological record shows that some ancient people abused their soils and that their civilizations were disrupted by the ecological and environmental consequences. One of history’s greatest problems has been the cultivation in and erosion of sloping lands. Repeated productivity losses on these soils would result in long-term costs to society, because this natural resource (capital) would eventually be depleted. Sustaining the long-term productivity and quality of all soil resources needs to be

1102

a national priority, which is essential to our nation’s prosperity. It is in the national interest that management strategies and techniques continue to be developed to monitor and sustain soil productivity and quality.

REFERENCES 1. Olson, G.W. Soils and the Environment—A Guide to Soil Surveys and their Applications; Chapman and Hall: New York, 1981. 2. Fowler, M.L. The Cahokia Atlas: A Historical Atlas of Cahokia Archaeology, Revised Ed.; University of Illinois at Urbana: Champaign, 1997. 3. Fowler, M.L. The Cahokia Atlas: A Historical Atlas of Cahokia Archaeology, Studies in Illinois Archaeology No. 6.; Illinois Historical Preservation Agency: Springfield, 1989. 4. Milner, G.R. The later prehistoric cahokia culture system of the Mississippi river valley: foundations, florescence, and fragmentation. J. World Prehist. 1990, 4, 1–43. 5. Pauketat, T.R. The Ascent of Chiefs: Cahokia and Mississippian Politics in Native North America; University of Alabama Press: Tuscaloosa, 1994. 6. Bareis, C.J.; Porter, J.W. American Bottom Archaeology: A Summary of the FAI-270 Project Contributions to the Culture History of the Mississippi River Valley; University of Illinois Press: Urbana, 1984.



7. Hall, R.L. Cahokia identity and interaction models of Cahokia Mississippian. In Cahokia and the Hinterlands: Middle Mississippian Cultures of the Midwest; Emerson, T.E., Lewis, R.B., Eds.; University of Illinois Press: Urbana, 1991. 8. Soil Survey Staff. Soil Taxonomy, a Basic System of Soil Classification for Making and Interpreting Soil Surveys; AH-436; U.S. Gov. Print. Office: Washington, DC, 1975. 9. Lowdermilk, W.C. Conquest of the Land through Seven Thousand Years, Agricultural Information Bulletin No. 99, 2nd Ed.; 1975. 10. Troeh, F.R.; Hobbs, J.A.; Donahue, R.L. Soil and Water Conservation for Productivity and Environmental Protection; Prentice-Hall: Englewood Cliffs, NJ, 1980. 11. Brown, L.R.; Wolf, E.C. Soil Erosion: Quiet Crisis in the World Economy; Worldwatch Institute: Washington, DC, 1984. 12. Lal, R. Soil degradation and the future of agriculture in sub-Saharan Africa. J. Soil Water Conserv. 1988, 43, 444–451. 13. Lee, L.K. Land use and soil loss: 1982 update. J. Soil Water Conserv. 1984, 39, 226–228. 14. USDA Soil Conservation Service Staff. Soil and Water Resources Conservation Act Program Report and Environmental Impact Statement, U.S. Govt. Print. Office: Washington, DC, 1981. 15. USDA Soil Conservation Service Staff. Preliminary Data, 1982. National Resources Inventory, Executive Summary: Washington, DC, 1983.

Mulch Farming Rattan Lal The Ohio State University, Columbus, Ohio, U.S.A.

INTRODUCTION Mulching, the practice of leaving crop residues and other biomass on the soil surface to protect it from climatic elements, is an ancient practice. Its application, however, is even more relevant now than ever before because of the serious global issues relevant to human population growth, scarcity of prime agricultural land, ecological benefits of nutrient recycling and optimizing rates of fertilizer input, severe problems of soil degradation, adverse water quality, and the accelerated greenhouse effect. It is an ecological approach to addressing the problem of sustainability, soil degradation environment quality, and nutrient cycling. Conventional agricultural practices, based on clean cultivation and intensive use of inorganic fertilizers and chemical amendments, make soils susceptible to degradative processes, e.g., erosion, acidification, and emission of greenhouse gases to the atmosphere. Stabilizing the atmospheric concentration of greenhouse gases is a major global challenge.[1] The global release of soil organic carbon (SOC) from agriculture is estimated at 800 Tg C=yr (T ¼ tera ¼ 1012 ¼ 1 million metric ton).[2] Soil restoration (for enhancement of SOC content) is an important and feasible option for mitigating the greenhouse effect,[3–5] and enhancing soil quality for increasing production and improving the environment.[6,7] These desirable properties and processes are achievable through frequent applications of crop residue mulch and other mulch farming techniques.

MULCH TYPES Mulch is any material, other than soil, specifically established at the soil–air interface to manage soil and water and create favorable environments for plant growth. Most mulches are organic materials of plant origin comprising crop residues, weed biomass, leaf litter, and by products of agroindustries, e.g., saw dust, rice husk, corn cobs, etc. However, there is a wide variety of mulch materials ranging from gravels and plastic to crop residues and planted fallows. Mulches can be broadly classified into two categories: organic and inorganic mulches (Fig. 1). 1. Organic mulches. Most mulch materials used in agriculture are of organic origin, and comprise crop residue used both ex situ and in situ. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042716 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

2. Inorganic mulches. Three most common inorganic mulches are plastic, gravels, and water.

MULCH PROCUREMENT FOR ARABLE LAND Crop Residues The most common practice of mulch procurement for use on agricultural land is the use of crop residues from the previous season’s crops. Crop residue is a renewable resource, and annually a large quantity of residues are produced in the world.[8,9] Residues produced in the world during 2001 wereare estimated at 3758 million Mg, of which 2802 million Mg (75%) are from cereals 305 million Mg (8%) from legumes, and 651 million Mg (17%) from other crops (Table 1). The amount of residue produced by seven crops in the tropics is estimated at 1838 million Mg (Table 2). Properly used as mulch, these residues can cover the entire arable land of 1500 million hectares at an average rate of 2.53 Mg=ha. Therefore crop residues are an important resource that need to be carefully and judiciously used for sustainable use of soil, water, and other natural resources. Crop residues management Management of crop residues is an important strategy of mulch procurement. Crop=plant residues, produced by annuals or perennials, in situ or ex situ, used as mulch can improve soil quality, increase productivity and sustainability, and enhance environmental quality, Fig. 2. Principal management alternatives outlined in Fig. 3 indicate four options: 1) mulching, 2) using as animal feed, 3) composting, and 4) burning. Using residues as mulch is the best option for controlling soil erosion, conserving soil water, and at the same time replenishing plant nutrient reserves. Planted Fallows and Cover Crops Cover crops, legumes, or grasses generally grown to produce biomass and provide ground cover, are a good source of mulch. Suitable cover crops are those that: 1) establish rapid ground cover; 2) improve soil fertility through biological nitrogen fixation and recycling 1103

1104

Mulch Farming

Fig. 1 Different types of mulch materials. (From Ref.[16].)

nutrients; 3) suppress weeds and other pests; and 4) produce a large quantity of biomass that can be used as mulch. Selected legumes and grasses, appropriate for the soil and environment, can ameliorate soil properties and produce the mulch material in a short time interval.[11] Deep rooted cover crops are also effective in improving subsoil properties.

BENEFITS OF MULCH FARMING Mulch farming is an ecological approach to agriculture. Properly used, it can improve and sustain agricultural production, enhance soil quality, and have minimal adverse effects on the environment (Fig. 4). Principal

benefits of mulch farming are briefly described in this figure. Soil-Surface Management and Erosion Control Soil erosion by water and wind is the principal degradative process,[12] and it adversely affects soil quality.[13,14] Mulching, at the rate of 4–6 Mg=ha, is an effective erosion-control measure.[15] Larson[16] recommended the use of crop residue mulch for erosion control in the mid-western U.S.A. Moldenhauer and Wischmeier[17] observed significant reduction in soil erosion by mulch farming practices. Both runoff and soil erosiondecrease

Table 2 Estimates of crop residues produced in the tropics Table 1 Estimates of crop residues produced by cereals, legumes, and other crops Residues produced (106 ton/yr) Crops

1991

2001

Cereals Legumes Oil crops Sugar crops Tubers Total

2563 238 162 340 145 3448

2802 305 108 373 170 3758

(From Ref.[9].)


Crop

Asia

Africa

South America

103 Mg Rice straw Rice husk Wheat Barley Sugarcane Cotton Oats Corn Total (From Ref.[10].)

772 154 380 34 54 6 2 166 1568

26 5 27 7 9 0.3 0.3 39 114

24 5 26 2 42 0.07 2 55 156

Mulch Farming

1105

Fig. 2 Interactive effects of mulch farming on soil quality=agronomic productivity, water quality, and greenhouse effect.

exponentially with increase in surface area covered by mulch (Fig. 5). Structural improvement by mulching is due to increase in SOC or humus content that facilitates the formation of organomineral complexes, and increase in soil biodiversity, notably the activity of earthworms, termites, and other beneficial soil fauna.[18]

Water Management There are four types of water: 1) surface water; 2) ground water; 3) soil water; and 4) rainwater. Mulching

affects all four types of water, but has a major impact on soil water and surface water. Mulching decreases runoff by improving infiltration rate, replenishes soil water by improving retention pores, and recharges groundwater by increasing percolation. Partitioning of rainwater into runoff, soil water, and groundwater components is influenced by mulching through its effect on the water balance. Mulching increases soil water and groundwater components by decreasing runoff and soil evaporation. Mulching improves soil water storage, and minimizes risks of drought. When the soil is wet at about field capacity level, as is the case soon

Fig. 3 Methods of procurement of mulch materials.


1106

Mulch Farming

Fig. 4 Economic, ecologic, and natural resource improvements by mulch farming.

after rainfall or irrigation, mulching decreases the evaporation rate and prolongs the duration during which soil remains moist (Fig. 6).

Soil Temperature Management Crop residue mulch has a buffering or moderating effect on soil temperature regime. It influences the extremes by decreasing the maximum and increasing

Fig. 5 A schematic depicting an exponential decline erosion with increasing soil cover by mulching.


the minimum soil temperature. Therefore, soil temperature range is narrower in mulched than unmulched soil. Mulching increases soil temperature in the morning and decreases it in the afternoon (Fig. 7). Plastic mulches are often used to increase soil temperature in horticultural and other high-value plants. Soil temperature effects of mulching are both due to direct and indirect effects. Direct effects are due to reduction in insolation or energy load on the soil surface. Mulching shades the soil and prevents the radiation reaching the soil surface. Indirectly,

Fig. 6 A schematic depicting mulch effects on soil.

Mulch Farming

1107

Table 4 Estimates of N contained in crop residue Crop

Asia

Africa and South America

World

3

10 Mg Rice Wheat Barley Sugarcane Cotton Oats Corn Total

4,862 1,899 222 226 64 15 781 8,069

328 308 60 212 4 12 788 1,712

5,346 5,651 1,521 526 69 325 5,393 18,832

(From Ref.[10].)

mulching influences soil temperature through its effect on soil’s heat capacity and thermal conductivity.

Tables 4, 5, and 6, respectively. Thus, cycling of nutrients contained in the residue is an important factor affecting soil quality and agronomic yields. It is feasible, therefore, that a considerable quantity of fertilizer can be saved by returning crop residues to the soil. Indeed, mulching and residues management are effective nutrient recycling mechanisms.

Nutrient Cycling and Soil Fertility Enhancement

Soil Biodiversity

Severe nutrient depletion in sub-Saharan Africa[19,20] and elsewhere can be controlled through nutrient cycling. Depending on plant species and management systems, crop residues contain a considerable amount of plant nutrients. Total amount of plant nutrients in crop residues range from 40 to 100 kg=Mg of residues.[8] On a global scale, crop residues contain 22.6 million Mg N, 3.6 million Mg P, 47.4 million Mg K (Table 3). Estimates of global fertilizer consumption show annual use of 113 million tons of NPK vs. 74 million Mg contained in crop residues. In comparison, annual fertilizer consumption is 16 million Mg of NPK in U.S. vis-a`-vis 9 million Mg contained in crop residues. Estimates of N, P, and K contained in crop residue in the world and in developing countries are shown in

Applications of crop residues mulch improve activity of soil micro- and macrofauna and biomass carbon. The relative magnitude of biomass carbon is a good indicator of soil quality. Mulching improves population and activity of soil fauna especially those of earthworms and termites. These organisms improve soil structure by burrowing, mixing and soil turnover activities. They decompose and mineralize crop residues and make nutrients available to plants.

Fig. 7 A schematic showing diurnal fluctuations in mulched and unmulched soil during summer.

Table 3 Nutrients contained in crop residues of cereals and legumes produced in the world Plant nutrients in residues (million tons/yr) Nutrient N P K Ca Mg N þ P þ K

Fertilizer use (million tons/yr)

U.S.A.

World

U.S.A.

World

2.98 0.47 5.70 1.87 0.85 9.15

22.62 3.58 47.39 12.11 6.16 73.59

10 2 4 — — 16

77 16 20 — — 113

(From Ref.[8].)


The Greenhouse Effect The atmospheric CO2 concentration has increased by approximately 30% from 280 ppm in 1850 to approximately 370 ppm in 2000.[21,22] This increase is attributed to two principal human activities: 1) land use

Table 5 Estimates of P contained in crop residue Crop

Asia

Africa and South America 103 Mg

Rice Wheat Barley Sugarcane Cotton Oats Corn Total (From Ref.[10].)

772 266 31 43 10 4 216 1342

53 37 8 40 1 3 149 291

1108

Mulch Farming

Table 6 Estimates of K contained in crop residue Crop

Asia

Africa and South America

World

3

10 Mg Rice Wheat Barley Sugarcane Cotton Oats Corn Total

4,600 2,355 235 361 64 40 1,230 8,885

446 418 73 338 4 32 1,013 2,324

5,322 7,758 2,374 839 69 852 6,824 24,038

(From Ref.[10].)

change and soil management (1.6 Pg C=yr; 1 Pg ¼ peta gram ¼ 1015 g) and 2) fossil fuel consumption (6.8 Pg C=yr). The rate of increase of atmospheric CO2 concentration is approximately 1.5 ppm or 3.5 Pg C=yr. Total crop residues production in the world estimated at 3460 million tons per annum[8] contains 1.5 Pg of C. This is approximately 43% of the annual atmospheric increase of 3.5 Pg=C, and cannot be ignored in the context of global C cycle. This C can have a major effect even if only 10% of the 1.5 Pg can be sequestered in soils as humus. In addition to the direct effect, mulching also sequesters C through erosion control, and increase in biomass production.[23]

MULCHING AND AGRICULTURAL SUSTAINABILITY Mulch farming affects agricultural sustainability through its impact on soil quality and productivity (Fig. 8). Mulch-induced changes in soil quality set-inmotion restorative processes leading to soil and water conservation, nutrient cycling and soil fertility improvement, improvement in soil structure and C sequestration, increase in soil biodiversity, and soil structure enhancement. Mulch farming improves energy use efficiency by 1) reducing fertilizer use through decreasing nutrient losses in runoff and erosion, and recycling plant nutrients; 2) decreasing irrigation needs through increasing infiltration, and reducing losses due to runoff and evaporation; 3) decreasing herbicide use through suppressing weed growth; and 4) decreasing fuel consumption through reduction of tillage operations and vehicular traffic. Crop residues are also a renewable source of energy that can be used in the farm household. Energy content of straw, depending on crop species and other factors, ranges from 3000 to 4000 kcal=kg.[24] Energy value of 1 Mg of residues may be as much as 9 million BTU,


Fig. 8 Mulch farming and agricultural sustainability.

18.6 109 J, or equivalent to two barrels of oil.[25] Estimates of fuel energy in crop residues show that cereal straw produced annually in the world is equivalent to 47 1018 J of energy (Table 7).[9] However, traditional use of biomass (crop residue, cattle dung, etc.) can have severe adverse impacts on the environment (Fig. 9).[26]

LIMITATION OF MULCH FARMING Biomass, the principal mulch material, has numerous alternative uses especially for resource-poor farmers of the tropics. Crop residues are used for feeding livestock for building fences and homesteads, and as fuel for cooking and household energy use. It is Table 7 Fuel energy of cereal straw produced in the world Parameter

Global value

Total Crop residue (million Mg=yr)

3758

Oil equivalent (106 barrels)

7560

Energy equivalent (Quads)

60

(From Ref.[8].)

Mulch Farming

1109

REFERENCES

Fig. 9 Biomass burning as household fuel and the environment.

a precious commodity, and little if any is left behind on the field as mulch on small farms in developing countries. Under some circumstances, mulches may also have negative effects on crop growth due to cool soil temperature during spring in temperate regions, high incidence of pests and pathogens, and N immobilization.

CONCLUSIONS Mulch farming maintains or enhances soil quality by conserving soil and water resources, improving SOC content, increasing soil biodiversity, strengthening nutrient cycling mechanisms, and regulating soil temperature regime. Mulch farming increases probability of achieving agricultural sustainability through: 1) increasing production; 2) decreasing costs of fertilizers and water inputs; and 3) preserving the resource base and its productive potential. Mulch farming has a potential to improve environmental quality by: 1) decreasing soil erosion and transport of chemicals in runoff, mulch farming decreases risks of eutrophication of surface waters and contamination of groundwater; and 2) reducing emission of radioactively active gases from soil to the atmosphere through increasing SOC content, stabilizing C within aggregates as organomineral complexes, decreasing losses due to soil erosion, and improving biomass production.


1. Engelman, R. Stabilizing the atmosphere: population, consumption, and greenhouse gases. In Population and Environment Program; Population Action International: Washington, DC, 1994; 48 pp. 2. Schlesinger, W.H. Evidence from chronosequence studies for a low carbon-storage potential of soils. Nature 1990, 348, 232–234. 3. Lal, R.; Kimble, J.M.; Levine, E.; Stewart, B.A., Eds. Soils and Global Change; CRC=Lewis Publishers: Boca Raton, 1995; 440 pp. 4. Lal, R.; Kimble, J.M.; Levine, E.; Stewart, B.A., Eds. Soil Management for Mitigating the Greenhouse Effect; CRC=Lewis Publishers: Boca Raton, FL, 1995; 385 pp. 5. Lal, R. Soil management and restoration for C sequestration to mitigate the accelerated greenhouse effect. Prog. Environ. Sci. 1999, 1, 307–326. 6. Doran, J.W.; Parkin, T.B. Defining and assessing soil quality. In Defining Soil Quality for a Sustainable Environment, SSSA Special Publication No. 35; SSSA: Madison, WI, 1994; 3–21. 7. Lal, R. Soil Quality and Agricultural Sustainability; Lal, R., Ed.; Ann Arbor Press: Chelsea, 1998; 3–11. 8. Lal, R. The role of residues management in sustainable agricultural systems. J. Sustain. Agric. 1995, 5, 51–78. 9. Lal, R. The world crop residue production and implications of its use as biofuel. Env. Intl. 2005, 31, 575–584. 10. Singh, Y.; Singh, B.; Timsina, J. Crop residue management for nutrient cycling and improving soil productivity in rice-based cropping systems in the tropics. Adv. Agron. 2005, 85, 269–407. 11. Smith, M.S.; Frye, W.B.; Varco, J.J. Legume winter cover crops. Adv. Soil Sci. 1987, 7, 95–140. 12. Oldeman, L.R. The global extent of soil degradation. In Soil Resilience and Sustainable Land Use; Greenland, D.J., Szabolcs, I., Eds.; CAB International: Wallingford, UK, 1994; 99–118. 13. Lal, R. Applying soil quality concepts for combating soil erosion. In Soil Quality and Soil Erosion; Lal, R., Ed.; CRC Press: Boca Raton, FL, 1999; 309–318. 14. Lal, R.; Mokma, D.; Lowery, B. Relation between soil quality and erosion. In Soil Quality and Soil Erosion; Lal, R., Ed.; CRC Press: Boca Raton, FL, 1999; 237–257. 15. Lal, R. Soil Erosion in the Topics: Principles and Management; McGraw-Hill: New York, 1990; 580 pp. 16. Larson, W.E. Protecting the soil resource base. J. Soil Water Conserv. 1981, 36, 13–16. 17. Moldenhauer, W.C.; Wischmeier, W.H. Soil and water losses and infiltration rates on ida silt loam as influenced by cropping systems, tillage practices, and rainfall characteristics. Soil Sci. Soc. Am. Proc. 1960, 24, 409– 413. 18. Lal, R. Soil structure and sustainability. J. Sustain. Agric. 1991, 1, 67–92. 19. Stoorvogel, J.J.; Scaling, E.M. Assessment of Soil Nutrient Depletion in Sub-Saharan Africa: 1983–2000;

1110

Report No. 28; The World Staring Center: Wageningen, The Netherlands, 1990. 20. Stoorvogel, J.J.; Smaling, E.M.A.; Janssen, B.H. Calculating soil nutrient balances in africa at different scales. I. Supranational Scale. Fertil. Res. 1993, 35, 227–235. 21. IPCC. In Technical Summary, Inter-Governmental Panel on Climate Change; WMO: Geneva, Switzerland, 1995; 44 pp. 22. Kennel, C. UC Revelle Program on Climate, Science and Policy; Scripps Institute of Oceanography: La Jolla, CA, 2000.


Mulch Farming

23. Lal, R. Residue management, conservation tillage and soil restoration for mitigating greenhouse effect by CO2 enrichment. Soil Tillage Res. 1997, 43, 81–107. 24. Stout, B.A. Handbook of Energy for World Agriculture; Elsevier: New York, 1990. 25. Anderson, A.W.; Anderson, J.F. On finding a use of straw. In Utilization and Recycles of Agricultural Wastes and Residues; Shuler, M.L., Ed.; CRC Press: Boca Raton, FL, 1980; 237–278. 26. Ramanathan, V.; Crutzen, P.J.; Kiehl, J.T.; Rosenfeld, D. Aerosols, climate and the hydrological cycle. Science 2002, 294, 2118–2124.

Mycorrhiza of Forest Ecosystems Ian A. Dickie University of Minnesota, St. Paul, Minnesota, U.S.A.

INTRODUCTION Most land plants (>80% of angiosperms and almost all gymnosperms) form mycorrhiza: symbiotic associations between plant roots and fungi that increase plant nutrient uptake.[1,2] Mycorrhizal fungi are an important component of the soil ecosystem, with ca. 200 m of mycorrhizal hyphae per gram of forest soil.[1] Several types of mycorrhiza occur, including arbuscular mycorrhiza (dominants of grasslands, many tropical and some temperate forests), ectomycorrhiza (dominants of boreal, most temperate and some tropical forests), and ericoid mycorrhiza (common in temperate and boreal forests and heathlands). Mycorrhiza plays a critical role in plant nutrient uptake, particularly of P, NH4þ, and organic nutrients. In return for increased nutrient acquisition, plants direct 10–30% of total fixed carbon to supporting mycorrhizas. This represents an important flow of carbon to the soil ecosystem, creating ‘‘mycorrhizosphere’’ communities in association with fungal hyphae.

MYCORRHIZAL TYPES Arbuscular mycorrhiza are the ancestral state of all land plants, and approximately 67% of all angiosperms and nearly all non-Pinaceae gymnosperms retain this association.[2] Arbuscular mycorrhiza are formed by most tropical and many temperate forest trees, as well as many forest-understory plants. Arbuscular mycorrhizas are characterized by internal colonization of the root, with the fungus forming characteristic structures (arbuscules) within the root cells. Arbuscular mycorrhizas are formed by fungi in the order Glomales. While less common than arbuscular mycorrhiza, ectomycorrhiza are particularly important in forest ecosystems (Fig. 1). Predominately ectomycorrhizal plant families include the Pinaceae, Fagaceae, Myrtaceae, Dipterocarpaceae, Salicaceae, Betulaceae, and the leguminous subfamily Caesalpinoideae, among others.[1] Although these species account for only around 3% of plant diversity, they include the dominants of most temperate and boreal forests and some tropical forests. In ectomycorrhiza, the fungus forms a mantle, which envelops root tips, and a Hartig net, Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120018262 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

which penetrates between the cells of the epidermis and cortex of the root but does not penetrate cells. The fungi forming ectomycorrhiza are primarily Basidiomycetes, as well as some Ascomycetes and some species of the Zygomycete genus Endogone. Many fungi forming ectomycorrhiza produce conspicuous sporocarps, some of which are economically important edibles (e.g., chanterelles, truffles, cepes, and matsutake). Ericoid mycorrhiza are exclusively formed by plants in the Ericales (Ericaceae, Epacridaceae, and Empetraceae). These plants are an important component of high altitude, temperate, and boreal forests and heathland ecosystems. Ericoid mycorrhiza form a loose weft of undifferentiated hyphae on the root surface, with extensive intercellular structures within root cells. Fungi forming ericoid mycorrhiza are primarily Ascomycetes, and some are Basidiomycetes.[3] In general terms, arbuscular mycorrhizal forests tend to dominate low-latitude and low-elevation sites with P-limited soils. Ectomycorrhizal forests become dominant with increasing latitude and elevation, and on more acidic and N limited soils. Ericoid mycorrhiza dominate at the highest elevations and latitudes, and on highly acid or organic soils.[1] Other forms of mycorrhiza also occur in forest ecosystems, including ectendomycorrhizas, arbutoid, orchid, and monotropoid mycorrhizas.[1] None of these types dominate forest ecosystems.

NUTRIENT UPTAKE Mycorrhiza can considerably increase plant nutrient uptake, and for many plants, they are the primary mechanism for nutrient uptake. All forms of mycorrhiza greatly increase the ability of plants to explore the soil. Estimates of the length of mycorrhizal hyphae in forest ecosystems range from 300 to 8000 m of fungal hyphae per meter root length.[4] By increasing effective root surface area, mycorrhizal fungi can substantially increase nutrient uptake, particularly of nutrients such as P and NH4þ that have low mobility in soils. The small size of fungal hyphae (down to 1 yr, the egg stage is followed by four juvenile stages before adulthood. In response to adverse environmental conditions (e.g., drought), nematodes may enter anabiosis, a state of metabolic dormancy that may last for years. Extensive information on nematode structure can be found in Ref.[9].

1116

BIOGEOGRAPHY Distribution and Abundance Soil nematodes are found in all biomes, with topsoil abundances ranging from 50,000=m2 in Antarctic deserts[10] up to tens of millions=m2 in temperate grasslands and deciduous forests.[11] Generally, nematodes occur wherever there is moisture and food available. Because these resources have heterogeneous distributions, nematode population are highly aggregated, which is a concern for quantitative soil sampling.[12] Active migration is limited in range, but nematodes may be dispersed over long distances (notably in anabiotic form) by wind, water, insects, birds, and mammals.[1]

Nematodes

with nematode-antagonistic crops, and physical methods such as tillage, steaming, and solarization.[20] Nematodes have numerous natural enemies including soil bacteria, fungi, predatory nematodes, and insects (Fig. 2),[21,22] but soil communities are not easily manipulated for biological control. Promising results have been obtained with Verticillium chlamydosporium, an egg parasite of cyst and root-knot nematodes.[23] The nematodetrapping fungus Arthrobotrys is commercially available, and controls nematodes in cultivated mushroom and tomatoes.[21]

BENEFICIAL NEMATODES Microbial-Feeding Nematodes

Species Diversity Less than 6000 terrestrial nematode species have been described, but it is estimated that there are least 1,00,000 species.[13] Species richness is generally greatest in temperate broadleaf forest (62 species per soil sample) followed by cultivated soil, grassland, tropical rainforest, temperate coniferous forest, and polar vegetation.[14] The structural and functional implications of nematode biodiversity are still being debated.[15]

HARMFUL NEMATODES Plant-parasitic nematodes can cause great damage to agricultural crops. Ectoparasitic nematodes stay outside the root and use their stylet to puncture root cells (e.g., thesting nematode Belonolaimus; the stunt nematode Tylenchorynchus). Endoparasitic nematodes enter the root and move around (the lesion nematode Pratylenchus) or stay in one feeding site (the root-knot nematode Meloidogyne; the cyst nematodes Globodera and Heterodera). The resulting root damage can cause wilting, stunting, nutrient deficiencies, and yield losses. Some species transmit viruses (Longidorus, Trichodorus, and Xiphinema). In other cases, nematodes may cause damage to above ground plant parts (the stem nematode Ditylenchus dipsaci, the wheat gall nematode Anguina tritici). Worldwide, nematode damage may amount to $78 billion annually.[16] Nematode problems of various crops and climates are reviewed by Refs.[17,18]. An important observation is that low levels of root herbivory could promote plant growth by stimulating root branching and exudation, which in turn enhances rhizosphere microbial activity and nutrient availability.[19] Nematode control may include sanitary measures, crop rotation, use of resistant cultivars, application of nematicides and organic amendments, intercropping


Microbial-feeding nematodes (bacterivores and fungivores) may promote plant growth by enhancing nutrient mineralization, notably of nitrogen.[24] As microbial-feeding nematodes have a higher carbonto-nitrogen ratio than microbes, and low production efficiency, they mineralize the nitrogen immobilized in microbial biomass. In addition, they can stimulate microbial activity (and thus nitrogen mineralization) by reducing microbial competition, enhancing oxygen diffusion, and transporting microbial propagules to new substrates. Using knowledge of microbivorous nematode population dynamics, the application of organic amendments can be timed to optimize nitrogen availability to crops.[25] Predacious and Omnivorous Nematodes By controlling microbivorous population, predacious (Fig. 2) and omnivorous nematodes prevent ‘‘overgrazing’’ of microbes and thereby indirectly control nutrient cycling processes.[26] In terms of biological control, it is unlikely that predacious nematodes could significantly contribute to plant-parasitic nematode control.[27] Their generation times are too long for a timely response to plant-parasite population peaks, and there is an evidence that predacious nematodes prefer microbial feeding, not plant-feeding, nematodes as diet. Predacious and omnivorous nematodes may be useful indicators of soil ecosystem health, because of their long life cycles and sensitivity to disturbance.[6] Entomopathogenic Nematodes Entomopathogenic (insect-parasitic) nematodes are successfully used in biological control of pest insects, including white grubs, leaf miners, and cockroaches.[28] Entomopathogenic nematodes (order Rhabditida)

Nematodes

1117

Fig. 2 A predacious nematode feeding on a gravid bacterivorous nematode. The predator itself is infected with Pasteuria penetrans (the ‘‘warts’’ attached to the predator’s cuticle are bacterial endospores).

carry a symbiotic bacterium. After penetrating the host insect, the nematode releases the bacterium, which kills the host within 72 hr, providing food to the nematode, which then multiplies inside the cadaver. Infective juveniles disperse to new hosts. (For detailed information on entomopathogenic nematodes, see Ref.[29].)

4.

5.

METHODS

6.

Methods for collection and extraction are reviewed by Refs.[17,30,31]. For ecological work with nematodes, useful methods are compiled by Ref.[32]. A bibliography of nematode identification keys is provided by Ref.[4], and an interactive diagnostic key to plant-parasitic, free-living, and predacious nematodes is found in Ref.[33].

7.

8.

REFERENCES 9. 1. Poinar, G.O., Jr. The Natural History of Nematodes; Prentice Hall: Englewood Cliffs, NJ, 1983. 2. Caenorhabditis elegans. http:==helios.bto.ed.ac.uk= mbx=fgn=wow=celegans.html (accessed August 2000). 3. Blaxter, M.L.; De Ley, P.; Garey, J.R.; Liu, L.X.; Scheldeman, P.; Vierstraete, A.; Vanfleteren, J.R.; Mackey, L.Y.; Dorris, M.; Frisse, L.M.; Vida, J.T.; Thomas,


10.

11.

W.K. A molecular evolutionary framework for the phylum Nematoda. Nature 1998, 392, 71–75. Bongers, T. Nematode identification literature. http:== www.spg.wau.nl=nema=ident_lit.html (accessed August 2000). Yeates, G.W.; Bongers, T.; De Goede, R.G.M.; Freckman, D.W.; Georgieva, S.S. Feeding habits in soil nematode families and genera—an outline for soil ecologists. J. Nematol. 1993, 25, 315–331. De Goede, R.G.M.; Bongers, T. The nematode maturity index. In Techniques in Nematode Ecology; Wheeler, T., Forge, T., Caswell-Chen, E., LaMondia, J.A., Eds.; http:==ianrwww.unl.edu=son=deGoede.htm (accessed August 2000) Society of Nematologists, 2000. Niles, R.K.; Freckman, D.W. From the ground up: nematode ecology in bioassessment and ecosystem health. In Plant and Nematode Interactions; Barker, K.R., Pederson, G.A., Windham, G.L., Eds.; Agronomy Monograph 36; American Society of Agronomy: Madison, WI, 1998; 65–85. Bongers, T. Maturity index literature. http:==www. spg.wau.nl=nema=MI_lit.htm (accessed August 2000). Bird, A.F.; Bird, J. The Structure of Nematodes; Academic Press: London, 1991. Freckman, D.W.; Virginia, R.A. Low-diversity Antarctic soil nematode communities: distribution and response to disturbance. Ecology 1997, 78, 363–369. Sohlenius, B. Abundance, biomass and contribution to energy flow by soil nematodes in terrestrial ecosystems. Oikos 1980, 34, 186–194.

1118

12. Webster, R.; Boag, B. Geostatistical analysis of cyst nematodes in soil. J. Soil Sci. 1992, 43, 583–595. 13. Andra´ssy, I. A short census of free-living nematodes. Fund. Appl. Nematol. 1992, 15, 187–188. 14. Boag, B.; Yeates, G.W. Soil nematode biodiversity in terrestrial ecosystems. Biodiv. Conserv. 1998, 7, 617–630. 15. Ettema, C.H. Soil nematode diversity: species coexistence and ecosystem function. J. Nematol. 1998, 30, 159–169. 16. Wheeler, T.; Forge, T.; Caswell-Chen, E.; LaMondia, J.A. Eds. Plant and Soil Nematodes: Societal Impact and Focus for the Future, A Report by the Committee on National Needs and Priorities in Nematology; Society of Nematologists: Lawrence, KS, 1994. 17. Nickle, W.R. Manual of Agricultural Nematology; Marcel Dekker: New York, 1991. 18. Luc, M.; Sikora, R.A.; Bridge, J. Plant-Parasitic Nematodes in Subtropical and Tropical Agriculture; CAB International: Wallingford, UK, 1990. 19. Denton, C.S.; Bardgett, R.D.; Cook, R.; Hobbs, P.J. Low amounts of root herbivory positively influence the rhizosphere microbial community in a temperate grassland soil. Soil Biol. Biochem. 1999, 31, 155–165. 20. NEMABASE—A database of the host status of plants to plant-parasitic nematodes. http:==www.ipm.ucdavis. edu=NEMABASE=index.html (accessed August 2000). 21. Poinar G.O., Jr.; Jansson, H.B. Diseases of Nematodes; CRC Press: Boca Raton, FL, 1988; Vol. I and II. 22. Natural enemies of nematodes: biological control images. http:==sacs.cpes.peachnet.edu=nemabc= (accessed August 2000). 23. Verticillium chlamydosporium and biological control of phytoparasitic nematodes. http:==www.area.ba.cnr. it=e085ac01=bkfair3444.html (accessed August 2000). 24. Freckman, D.W. Bacterivorous nematodes and organicmatter decomposition. Agric. Ecosyst. Environ. 1988, 24, 195–217.


Nematodes

25. Ferris, H.; Venette, R.C.; Lau, S.S. Dynamics of nematode communities in tomatoes grown in conventional and organic farming systems, and their impact on soil fertility. Appl. Soil Ecol. 1996, 3, 161–175. 26. Hunt, H.W.; Coleman, D.C.; Ingham, E.R.; Ingham, R.E.; Elliott, E.T.; Moore, J.C.; Rose, S.L.; Reid, C.P.P.; Morley, C.R. The detrital food web in a short grass prairie. Biol. Fert. Soils 1987, 3, 57–68. 27. Yeates, G.W.; Wardle, D.A. Nematodes as predators and prey: relationships to biological control and soil processes. Pedobiologia 1996, 40, 43–50. 28. Smart G.C., Jr. Entomopathogenic nematodes for the biological control of insects. J. Nematol. 1995, 4-S, 529–534. 29. Gaugler, R.; Kaya, H. Entomopathogenic Nematodes in Biological Control; CRC Press: Boca Raton, 1990. 30. Ingham, R.E. Nematodes. In Methods of Soil Analysis. Part 2—Microbiological and Biochemical Properties; Bigham, J.M., Ed.; Soil Science Society of America: Madison, WI, 1994; 459–490. 31. McSorley, R. Sampling and extraction techniques for nematodes. In Techniques in Nematode Ecology; Wheeler, T., Forge, T., Caswell-Chen, E., LaMondia, J.A., Eds.; http:==ianrwww.unl.edu=son=McSorley.htm (accessed August 2000) Society of Nematologists, 2000. 32. Wheeler, T.; Forge, T.; Caswell-Chen, E.; LaMondia, J.A., Eds. Techniques in Nematode Ecology; http:== ianrwww.unl.edu=son=Ecology_Manual_TOC.htm (accessed August 2000) Society of Nematologists, 2000. 33. Interactive diagnostic key to plant parasitic, freeliving and predacious nematodes. http:==ianrwww.unl.edu= ianr=plntpath=nematode=key=nemakey.htm. 34. Golden, A.M. Classification of the genera and higher categories of the order Tylenchida (Nematoda). In Plant Parasitic Nematodes; Zuckerman, B.M., Mai, W.F., Rohde, R.A., Eds.; 1971; Vol. 1, 191–232.

Nitrate Leaching Index Harold M. van Es Cornell University, Ithaca, New York, U.S.A.

Jorge A. Delgado Soil Plant Nutrient Research Unit, United States Department of Agriculture-Agricultural Research Service (USDA-ARS), Fort Collins, Colorado, U.S.A.

INTRODUCTION Nitrogen is an essential plant nutrient and key to the sustainability and economical viability of agricultural systems. On average, current N use efficiencies (NUE) are being reported to be about 50%, and the economic worldwide average N losses are equivalent to millions of U.S. dollars. The Environmental Protection Agency considers water with over 10 mg NO3-N L1 concentration unsafe for drinking purposes,[2] and several studies have shown that nitrate leaching from agricultural systems can readily cause this level to be exceeded.[3,4] The Nitrate Leaching Index (NLI) is an indicator of the potential for nitrate to reach shallow groundwater and incorporates information on soils and precipitation.[5] It can be used as a tool to identify land areas where good N management is critical. This article discusses the processes affecting nitrate leaching in soils, explains the current nitrate leaching index and its strengths and weaknesses, and provides suggestions for improvement.

PROCESSES The pathways of N losses need to be viewed within the context of the N cycle and the need to budget all N sources.[6] Nitrogen can readily be transformed from organic forms and ammonia fertilizer forms into NO3-N, which is mobile in soil and therefore subject to leaching. However, plants require most N to be in the NO3 form; therefore the main objective of good agricultural N management is to have sufficient NO3-N available in the root zone for adequate plant nutrition, but to avoid any excess N that would lead to high leaching losses. This is a challenge in many crop production systems because 1) soil by its very nature is ‘‘leaky’’ and water percolation and leaching of chemical substances is often unavoidable; 2) N additions cannot always be precisely managed, especially when dealing with organic sources; and 3) soil N is subject to other loss pathways such as denitrification and subsequent gaseous emission, ammonia volatilizaEncyclopedia of Soil Science DOI: 10.1081/E-ESS-120026542 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

tion, surface runoff, and erosion. These processes are often difficult to predict because they are affected by site-specific factors (e.g., fine-textured soils experience more denitrification than coarse-textured ones[7]) and are strongly dependent on spurious climate factors such as temperature and precipitation. Improved management of N requires an understanding of these processes within the context of the N cycle[6] and a recognition that inadequate predictability of the various loss pathways often causes farmers to use excess (insurance) N rates to avoid plant N deficiencies under worst-case scenarios. Nevertheless, agricultural N can be effectively managed to reduce NO3-N leaching,[8] and the NLI is a tool developed for this purpose. We will discuss the current tool and its applications, as well as possible future improvements.

THE N LEACHING INDEX The extent of soil water percolation and thereby NO3-N leaching potential mostly depends on permeability, pore size distribution, soil depth to a restrictive layer, artificial drainage, and precipitation amount and distribution over the year. For a given precipitation pattern, welldrained soils generally have greater N leaching potential than poorly drained soils. The NLI estimates this percolation potential from the Natural Resources Conservation Service (NRCS) soil hydrologic unit designations, which range from Group A soils with high percolation rates (generally coarse-textured and well-drained) to Group D soils with low percolation rates (generally fine-textured and poorly drained). The NLI is an indicator of the potential for nitrate to reach shallow groundwater and incorporates information on soils and precipitation:[5] NLI ¼ Percolation Index Seasonal Index The Percolation Index (PI) is a function of the annual average precipitation (PA, in inches) and soil hydrologic group, and is estimated by the following 1119

1120

equations for each of the hydrologic groups:[9] Hydrologic Group A: PI ¼ ðPA 10:28Þ2 =ðPA þ 15:43Þ Hydrologic Group B: PI ¼ ðPA 15:05Þ2 =ðPA þ 22:57Þ Hydrologic Group C: PI ¼ ðPA 19:53Þ2 =ðPA þ 29:29Þ Hydrologic Group D: PI ¼ ðPA 22:67Þ2 =ðPA þ 34:00Þ Soils with mixed hydrologic grouping, e.g., B=D, as a result of artificial drainage, generally use the higher percolation potential in the NLI assessment. The Seasonal Index (SI) is determined by the annual precipitation (PA) and the sum of the nongrowing season precipitation (typically the winter period in the United States, PN in inches): SI ¼ ð2 PN =PA Þ1=3 The SI reflects the fact that most water percolation occurs from precipitation after crop senescence and removal, as a result of the consequent drop in transpiration rate. Nitrate Leaching Index values are generally interpreted as follows: NLI < 2: Low nitrate leaching risk 2 < NLI < 10: Medium nitrate leaching risk NLI > 10: High nitrate leaching risk Despite its simplicity, the NLI is generally effective in identifying high-leaching scenarios from low-leaching scenarios[10] and has been widely adopted in the United States for nutrient management planning on farms.

FRAMEWORK FOR AN IMPROVED N LEACHING INDEX Although the current NLI allows for the assessment of NO3-N leaching potential, it also has significant limitations because of its oversimplification of complex processes.[11] The NLI does not actually estimate the leaching of NO3-N and therefore cannot be directly tied to an enforceable water-quality standard (e.g., 10 mg L1 NO3-N). The index also ignores important processes that affect N leaching, such as denitrification, which is especially significant in fine-textured soils.[4] van Es, czymmek, and ketterings.[10] concluded that despite this shortcoming, the NLI still effectively identifies soil and climate regions of high and low leaching potential.


Nitrate Leaching Index

Another major limitation of the NLI is the lack of a connection with management practices, such as N application rates and timing, crop type, and crop rotations, which greatly influence NO3-N leaching. This concern has, in some cases, been addressed by providing specific management guidelines for fields with high NLI values, such as conservative N applications, optimum timing of applications, cover cropping, etc.[10] Shaffer and Delgado[11] proposed a framework for an improved NLI through a technologically based index that requires the integration of databases describing soil hydraulic properties, climate, off-site factors and management, the use of simulation models, and the internet. This NLI is proposed to use a dynamic ranking system similar to that for the P Index,[12] which can facilitate a preliminary assessment by classifying combinations of management scenarios, soil conditions, climate, crop types and off-site factors into leaching potential ratings. This NLI should be simple enough that it can be used by consultants, agronomists, conservationists, farmers, and technical personnel. It is proposed to have a tiered approach, where initial efforts involve simple screening to separate the higher NO3-N leaching potential scenarios from the lower ones.[11] Higher tiers of this NLI should be capable of assessing N leaching potential within the context of more complex N dynamics, transformations, and mobility in the root zone. Also, this NLI must be developed so it can be simultaneously used with the P Index to perform integrated analyses of nutrient loss potential, and allows for the evaluation of environmental tradeoffs associated with management practices. For example, early-fall manure application in the cool temperate regions of the United States is often optimal for avoiding P runoff losses, but poses high N leaching risk compared to other seasonal application periods.[10] At the first tier of application, the NLI should be simple enough to be applied across large regions. At the higher, more complex levels, the NLI should be capable of evaluating the leaching potential related to site-specific management scenarios by being coupled to GIS and capable of identifying variable NO3-N leaching areas across single fields of various soil characteristics and management scenarios. A final attribute of a new NLI should be its national scope and consistency, but not necessarily identical formulation. This allows for uniform standards and effective communication among technical service providers and scientists.[11] A new NLI needs to be developed with mechanistic dynamic simulation models (e.g., GPFARM, EPIC, LEACHM, NLEAP, GLEAMS, RZWQM) that can account for the complex N pathways, transformations, and interactions with other nutrients.[13] This is especially important for the evaluation of cases where organic N sources are being applied to the fields, which are generally of greatest concern with nutrient losses.

Nitrate Leaching Index

CONCLUSIONS The current NLI is a rapid assessment tool that evaluates the N leaching potential based on basic soil and climate information. It is the basis for many current nutrient management planning efforts, but has considerable limitations because of 1) an oversimplification of the processes affecting N leaching, and 2) a lack of management considerations. Improved N management in the landscape requires a new NLI that considers the complex interactions of climate conditions, soil characteristics, crop type, off-site factors, and management scenarios. A tiered approach is proposed to achieve these multiple objectives.[11] REFERENCES 1. Hallberg, G.R. Nitrate in ground water in the United States. In Nitrogen Management and Ground Water Protection; Follett, R.F., Ed.; Elsevier: New York, NY, 1989; 35–74. Chapter 3. 2. USEPA (United States Environmental Protection Agency). Federal Register: Washington, DC, May 1989; 2254 FR 22062. 3. Randall, G.W.; Huggins, D.R.; Russelle, M.P.; Fuchs, D.J.; Nelson, W.W.; Anderson, J.L. Nitrate losses through subsurface tile drainage in conservation reserve program, alfalfa and row crop systems. J. Environ. Qual. 1997, 26, 1240–1247. 4. Sogbedji, J.M.; van Es, H.M.; Yang, C.L.; Geohring, L.D.; Magdoff, F.R. Nitrate leaching and N budget as affected by maize N fertilizer rate and soil type. J. Environ. Qual. 2000, 29, 1813–1820.


1121

5. Williams, J.R.; Kissel, D.E. Water percolation: an indicator of nitrogen-leaching potential. In Managing Nitrogen for Groundwater Quality and Farm Profitability; Follet, R.F., Ed.; Soil Science Society of America Inc.: Madison, WI, 1991; 59–83. 6. Delgado, J.A. Quantifying the loss mechanisms of nitrogen. J. Soil Water Conserv. 2002, 57, 389–398. 7. Moisier, A.R.; Doran, J.W.; Freney, J.R. Managing soil denitrification. J. Soil Water Conserv. 2002, 57, 505–513. 8. Meisinger, J.J.; Delgado, J.A. Principles for managing nitrogen leaching. J. Soil Water Conserv. 2002, 57, 485–498. 9. Pierce, F.J.; Shaffer, M.J.; Halvorson, A.D. Screening procedure for estimating potentially leachable nitrate– nitrogen below the root zone. In Managing Nitrogen for Groundwater Quality and Farm Profitability; Follet, R.F., Ed.; Soil Science Society of America, Inc.: Madison, WI, 1991; 259–283. 10. van Es, H.M.; Czymmek, K.J.; Ketterings, Q.M. Management effects on N leaching and guidelines for an N leaching index in New York. J. Soil Water Conserv. 2002, 57, 499–504. 11. Shaffer, M.J.; Delgado, J.A. Essentials of a national nitrate leaching index assessment tool. J. Soil Water Conserv. 2002, 57, 327–335. 12. Sharpley, A.N.; Daniel, T.; Sims, T.; Lemunyon, J.; Stevens, R.; Parry, R. Agricultural Phosphorus and Eutrophication; USDA-ARS, 1999; ARS-149. 37 pp. 13. Delgado, J.A. Use of simulations for evaluation of best management practices on irrigated cropping systems. In Modeling Carbon and Nitrogen Dynamics for Soil Management; Shaffer, M.J., Ma, L., Hansen, S., Eds.; Lewis Publishers: Boca Raton, Florida, 2001; 355–381.

Nitrate Leaching Management John J. Meisinger United States Department of Agriculture (USDA), Beltsville, Maryland, U.S.A.

Jorge A. Delgado United States Department of Agriculture (USDA), Fort Collins, Colorado, U.S.A.

Ashok Alva United States Department of Agriculture (USDA), Prosser, Washington, U.S.A.

INTRODUCTION Nitrate leaching occurs when the soil nitrate–nitrogen (NO3–N) concentrations are high and water moves beyond the root zone. Leaching losses in modern agriculture commonly account for 10–30% of the nitrogen (N) additions.[1–3] Leaching can contribute to nitrate enrichment of groundwater, which has a health advisory limit of 10 mg NO3–N=L, and to eutrophication of surface waters that can lead to the development of hypoxic zones in receiving waters. Managing leaching requires development of site-specific practices that should be based on an: understanding of the soil–crophydrologic cycle, avoiding excess N by applying a N rate to meet expected yields, and applying N in phase with crop demand. Specific nitrate management approaches include adjusting irrigation inputs according to site water needs, employing cropping systems that fully utilize soil-water resources, and utilizing within-season and real-time N monitoring tools. The goal of this entry is to discuss practical techniques to reduce nitrate leaching from modern agriculture.

APPROACHES FOR DECREASING NITRATE LEACHING Primary Techniques for Managing Nitrate Leaching The major nitrate leaching management techniques include understanding the soil–crop-hydrologic cycle, applying the proper rate of N, and applying N in phase with crop demand. Because water movement is the driving force for nitrate leaching, it is essential to understand the hydrologic cycle before developing site-specific leaching management strategies.[1] Understanding the hydrologic cycle of the site should identify the most likely times when leaching can occur, i.e., the times when the soil is at near field capacity and water inputs exceed water use by evapotransporation (ET). 1122 Copyright © 2006 by Taylor & Francis

In humid regions, leaching is usually minimal in summer when ET exceeds precipitation. But during the winter-to-spring season, precipitation exceeds ET and leaching is common, as shown by the percolate data of Fig. 1, adapted from large monolith lysimeter data from Ohio, U.S.A.[3] In irrigated agriculture leaching is most likely to occur during the growing season from excess water inputs on coarse-textured soils. Controlling leaching in humid regions should focus on management practices to keep the soil NO3-N levels low during the fall season and on providing a crop N sink during the nongrowing season, such as a grass cover crop. In irrigated agriculture, avoiding excess irrigation inputs through irrigation scheduling is an effective method to manage leaching. Avoiding excess N inputs, from fertilizer and=or manure, is the most fundamental approach to control leaching because crop N use efficiency is usually high for plants responding to N. By contrast, N rates on the nonresponsive part of the yield curve are associated with lower efficiencies and high levels of unused N that is vulnerable to leaching. The data of wheat grain N removals (assuming 20 kg N=t grain) from the Rothemstad Broadbalk Study is summarized in Fig. 2; the associated estimate of N leaching is adapted from Gouldings.[4] These data show only small leaching losses for N rates in the responsive part of the curve, with a marked increase in leaching in the nonresponsive portion of the curve (Fig. 2). Applying the proper N rate usually involves estimating the crop N need from the expected yield and then subtracting N credits from residual NO3-N, irrigation N inputs, prior legume crop credits, and adjusting for available manure N.[1] Applying N in phase with crop demand is also a fundamental approach for managing leaching that involves applying N after crop establishment, or after each forage harvest in grass-forage systems. Timing N applications to closely match crop N uptake minimizes the time that N is exposed to uncontrolled rainfall inputs and thus leads to lower leaching losses. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120026541 Copyright # 2006 by Taylor & Francis. All rights reserved.

Nitrate Leaching Management

Fig. 1 Monthly lysimeter percolate (upper panel) and N leached (lower panel, as percent of total yearly N leached) for large monolith lysimeters in Coshocton, Ohio, U.S.A. (From Ref.[3].)

Irrigation and Cropping System Techniques for Managing Nitrate Leaching Other nitrate management techniques include irrigation scheduling, modifying the cropping system, and developing riparian zones and conservation acres. Irrigation water management techniques are based on irrigation scheduling, which adds irrigation in accordance with ET and soil water status. Irrigation scheduling can reduce leaching by avoiding excess water applications, by adjusting water inputs to local weather, and by adding water based on local soil properties and local soil water content. For example, irrigating at 85% of ET, compared to 100% of ET, reduced leaching losses from about 110 to 60 kg N=ha on a sandy soil in Nebraska, U.S.A.[5] Similarly,

1123

Diez et al.[6] reported that scheduled irrigations to account for crop water demands can reduce the leaching of water and of NO3-N by four times, compared to excess irrigation. Meisinger and Delgado[1] summarized cropping system effects on leaching and concluded that adding a grain legume, such as soybeans, can partially reduce leaching compared to continuous corn (typically 5–10%), but adding a forage crop such as alfalfa or a forage grass, can substantially reduce leaching (typically 70–90%). However, changing the cropping system requires simultaneous changes in marketing and=or livestock enterprises. Delgado[7] reported that cropping systems that have shallowrooted crops and are heavily fertilized are most susceptible to NO3-N leaching. However, N recovery can be significantly improved by adding a deep-rooted crop such as malting barley. The barley served as a scavenger crop and even mined NO3-N from the system.[7] Adding a grass cover crop is another cropping system approach to reduce leaching by converting mobile NO3-N to immobile plant protein. A review of the effects of cover crops on water quality[8] concluded that grass covers can reduce nitrate leaching by an average of 70% compared to no-cover, while legume covers can reduce leaching by about 20%. Cover crops are especially useful in humid regions, where the winter leaching potential is high (Fig. 1). Developing riparian zones and conservation reserve areas can reduce the N loss from production fields into adjacent streams, especially in tile-drained watersheds.[1]

Other Approaches for Managing Nitrate Leaching Meisinger and Delgado[1] have also discussed other leaching management techniques. Nitrification inhibitors can delay leaching losses but are most effective in conjunction with other techniques, such as reduced N rates. Changing tillage practices usually have only secondary effects on leaching. Drainage-ditch water control measures can reduce nitrate transport to streams by impounding water and promoting denitrification, but have limited geographic applicability. Improved N application equipment is a direct approach to improve N application accuracy and avoiding excess N rates; new application equipment can apply N at the meter and submeter spatial scale.

Within-Season Monitoring Techniques for Managing Nitrate Leaching Fig. 2 Wheat grain N yield and N leaching for various fertilizer rates; Broadbalk Study, Rothamsted, U.K. (From Ref.[4].)


In-season testing and real-time sensors are recently developed tools for managing leaching. These monitoring approaches seek to identify N sufficient or deficient

1124

areas within a field and adjust subsequent N applications accordingly. The leaf chlorophyll meter (LCM) compares leaf greenness to the LCM reading in a well-fertilized reference strip. The relative LCM value has been shown to successfully predict the need for extra fertilizer N, especially in irrigated systems where N can be applied with irrigation water.[9] The presidedress soil nitrate test (PSNT) measures NO3-N in the surface 30 cm of soil when corn is about 30 cm tall, and compares it to a sufficiency concentration of 20–25 mg NO3-N=kg. The PSNT is most useful for diagnosing N sufficient sites. A Connecticut, U.S.A., study[10] compared conventional N management with the PSNT and reported N leaching loses of 50 and 20 kg N=ha, respectively. Plant stem or petiole NO3-N tests measure plant NO3-N sufficiency at specific growth stages and are commonly used in vegetable crops.[1] Real-time sensors that measure crop biomass and greenness based on remotely sensed or tractormounted units, sense red and near infra-red reflectance,[1] and are new tools for managing N leaching. Data from real-time sensors have a meter to submeter resolution and can be combined with geographically mapped data of soil properties and previous crop yields to produce a real-time assessment of N status, potential yield, and the suggested N application based on crop simulation models such as nitrate leaching and economic analysis package (NLEAP).[7] These above seasonal and real-time monitors have been shown to reduce nitrate leaching by identifying N sufficient areas and avoiding applications of excess N; they can also increase profitability by identifying N deficient areas with excellent small-scale resolution. An example of the benefits from Precision Agriculture has been reported by Bausch and Delgado[11] who reported that using GIS and remote sensing tools for in-season N management resulted in N applications that required 52% less N than that used under conventional practices (214 kg N=ha=yr). On average, the in-season N management saved 102 kg N=ha=yr which was worth about $55=ha=yr, without yield reductions during two consecutive growing seasons. These results show that precision management can significantly improve N efficiency of corn systems without reducing grain yields, thus minimizing the potential for NO3-N leaching.

CONCLUSIONS Nitrate leaching is a significant N loss process for agriculture that must be managed to minimize nitrate enrichment of groundwater and surface waters. Managing nitrate leaching should involve application of the basic principles of understanding the site’s hydrologic cycle, avoiding excess rates of N, and applying N in phase with crop demand. Other specific techniques to


Nitrate Leaching Management

reduce leaching include use of irrigation scheduling, grass cover crops, within-season monitoring with soil NO3-N tests or leaf chlorophyll meters, real-time sensors using red and near infra-red reflectance, and use of remote sensing with geographic information systems and simulation models to identify the best combination of practices to control leaching. Application of a specific combination of the above practices should increase crop N recovery with concomitant reductions in NO3-N leaching.

REFERENCES 1. Meisinger, J.J.; Delgado, J.A. Principles for managing nitrogen leaching. J. Soil Water Conserv. 2002, 57 (6), 485–498. 2. Legg, J.O.; Meisinger, J.J. Soil nitrogen budgets. In Nitrogen in Agricultural Soils; Stevenson, F.J., Bremner, J.M., Hauck, R.D., Keeney, D.R., Eds.; American Society of Agronomy: Madison, WI, 1982; Monograph No. 22, 503–566. 3. Chichester, F.W.; Smith, S.J. Disposition of 15N-labeled fertilizer nitrate applied during corn culture in field lysimeters. J. Environ. Quality 1978, 7 (2), 227–233. 4. Gouldings, K. Nitrate leaching from arable and horticultural land. Soil Use Manage. 2000, 16 (1), 145–151. 5. Hergert, G.W. Nitrate leaching through sandy soil as affected by sprinkler irrigation management. J. Environ. Quality 1986, 15 (3), 272–278. 6. Diez, J.A.; Roman, R.; Caballero, R.; Caballero, A. Nitrate leaching from soils under a maize-wheat-maize sequence, two irrigation schedules and three types of fertilizers. Agric. Ecosyst. Environ. 1997, 65 (1), 189–199. 7. Delgado, J.A. Use of simulations for evaluation of best management practices on irrigated cropping systems. In Modeling Carbon and Nitrogen Dynamics for Soil Management; Shaffer, M.J., Ma, L., Hansen, S., Eds.; Lewis Publishers: Boca Raton, FL, 2001; 355–381. 8. Meisinger, J.J.; Hargrove, W.L.; Mikkelsen, R.B.; Williams, J.R.; Benson, V.W. Effect of cover crops on groundwater quality. In Cover Crops for Clean Water; Hargrove, W.L., Ed.; Soil and Water Conservation Society of America: Ankeny, IA, 1991; 57–68. 9. Blackmer, T.M.; Schepers, J.S. Use of a chlorophyll meter to monitor nitrogen status and schedule fertigation for corn. J. Production Agric. 1995, 8 (1), 56–60. 10. Guillard, K.; Morris, T.F.; Kopp, K.L. The pre-sidedress soil nitrate test and nitrate leaching from corn. J. Environ. Quality 1999, 28 (6), 1845–1852. 11. Bausch, W.C.; Delgado, J.A. Ground base sensing of plant nitrogen status in irrigated corn to improve nitrogen management. In Digital Imaging and Spectral Techniques: Applications to Precision Agriculture and Crop Physiology ; Van Toai, T., Major, D., McDonald, M., Schepers, J., Tarpley, L., Eds.; Am. Soc. Agronomy Spec. Pub. 66: Madison, WI, 2003; 151–163.

Nitrogen and Its Transformations Oswald Van Cleemput Pascal Boeckx Ghent University, Ghent, Belgium

INTRODUCTION Nitrogen (N) is essential to all life. It is the nutrient that most often limits biological activity. In agricultural and natural ecosystems, N occurs in many forms covering a range of valence states from 3 to þ5. The change from one valence state to another depends primarily on environmental conditions. The transformations and flow from one form to another constitute the basics of the soil N cycle (Fig. 1). The use of N fertilizers has become essential to increase the productivity of agriculture, and has resulted in an almost doubling of the global food production in the past 50 years. However, this also implies that the natural N cycle has substantially been disturbed. In the following paragraphs an overview of the different N transformation processes in the soil is given.

THE NITROGEN CYCLE: GENERAL Atmospheric N2 gas (valence 0) can be converted by lightening to various oxides and finally to nitrate (NO3) (valence þ5), which can be deposited and taken up by growing plants. Also N2 gas can be converted to ammonia (NH3, valence 3) by biological N2 fixation, with the NH3 participating in a number of biochemical reactions in the plant. When plant residues decompose the N-compounds undergo a series of microbial conversions (mineralisation) leading first to the formation of ammonium (NH4þ) (valence 3) and possibly ending up in NO3 (nitrification). Under anaerobic conditions NO3 can be converted to various N-oxides and finally to N2 gas (denitrification). When mineral or organic N fertilizers are used they also undergo the same transformation processes and influence the rate of other Ntransformations. In considering the soil compartment, there can be N gains (such as biological N2 fixation) as well as N losses (such as leaching and denitrification). Furthermore N can be exported from the soil via harvest products, or immobilized in soil organic matter.

NITROGEN TRANSFORMATIONS IN THE SOIL The principal forms of N in the soil are NH4þ, NO3 or organic N-substances. At any moment, inorganic Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001574 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

N in the soil is only a small fraction of the total soil N. Most of the N in a surface soil is present as organic N. It consists of proteins (20–40%), amino sugars, such as the hexosamines (5–10%), purine and pyrimidime derivates (1% or less), and complex unidentified compounds formed by reaction of NH4þ with lignin, polymerization of quinones with N compounds and condensation of sugars and amines. In the subsoil, an important fraction of the present N can be trapped in clay lattices (especially illitic clays) as nonexchangeable NH4þ and is consequently largely unavailable. Organic substances slowly mineralize by microorganisms to NH4þ, which could be converted by other microorganisms to NO3 (see further). The NH4þ can be adsorbed to negatively charged sites of clay minerals and organic compounds. This reduces its mobility in the soil compared to the more mobile NO3 ion. Microorganisms can use both NH4þ and NO3 to satisfy their need for N. This type of N transformation is called microbial immobilization. The ratio between carbon (C) and N (C : N ratio) in organic matter determines whether immobilization or mineralization is likely to occur. When utilizing organic matter with a low N content, the microorganisms need additional N, decreasing the mineral N pool of the soil. Thus, incorporation of organic matter with a high C : N ratio (e.g., cereal straw) results in immobilization. Incorporation of organic matter with a low C : N ratio (e.g., vegetable or legume residues) results in N-mineralization. A value of the C : N ratio of 25 to 30 is often taken as the critical point toward either immobilization or mineralization. Nitrification is a two-step process. In the first step NH4þ is converted to nitrite (NO2) (valence þ3) by a group of obligate autotrophic bacteria known as Nitrosomonas species. The second step is carried out by another group of obligate autotrophic bacteria known as Nitrobacter species. Also a few heterotrophs can carry out nitrification, usually at much lower rates. Soil water and aeration are crucial factors for nitrification. At a water potential of 0 kPa (saturation), there is little air in the soil and nitrification stops, due to oxygen limitation; nitrification is greatest near field capacity (33 kPa in medium- to heavy-textured soils, to 0 to 10 kPa in light sandy soils). Also in 1125

1126

Nitrogen and Its Transformations

Fig. 1 The soil N cycle. The white compartment represents the atmosphere; the light gray compartment represents the biosphere and the dark gray compartment the subsoil.

dry soils NH4þ and sometimes NO2 accumulate presumably because Nitrobacter species are more sensitive to water stress than the other microorganisms. Nitrification is slow in acid conditions with an increasing rate at increasing pH. Mainly under alkaline conditions, nitrite is also accumulating, because Nitrobacter is known to be inhibited by ammonia, which is formed under alkaline conditions. Nitrification is a process that acidifies the soil as protons (Hþ) are liberated: NH4þ þ 2O2 ! NO3 þ 2Hþ þ H2 O During nitrification minor amounts of nitrous oxide (N2O) (valence þ1) and nitric oxide (NO) (valence þ2) are formed. Both compounds have environmental consequences, discussed below. The effect of temperature on nitrification is climate dependent. There is a climatic selection of species of nitrifiers, with those from cooler regions having lower temperature optima and less heat tolerance than species from warmer regions. All above-mentioned factors influencing nitrification also influence the nitrifying population. The population and activity of nitrifiers can be reduced by the use of nitrification inhibitors, such as dicyanodiamide, nitrapirin and neem (Azadirachta indica) seed cake. They are used mostly to retard


the nitrification of manure; otherwise their practicality is controversial and they are not extensively used. More details about nitrification and nitrification inhibitors can be found in McCarty[1] and Prosser.[2] NITROGEN INPUT PROCESSES Atmospheric Nitrogen Deposition The total atmospheric N (NH4þ and NO3) deposition is in the order of 10–40 kg N ha1 yr1 in much of north-western and central Europe and some regions in North America. It ranges from 3–5 kg N ha1 yr1 in pristine areas.[3] It is originating from previously emitted NH3 and NOx from agricultural and industrial activities or traffic. Biological Nitrogen Fixation Rhizobium species living in symbiotic relationship in root nodules of legumes, e.g. clover (Trifolium), lucerne (Medicago), peas (Pisum) and beans (Faba)— can convert atmospheric N2 gas to NH3, which is further converted to amino acids and proteins. Parallel to this process, the rhizobium species receive from the legume the energy they need to grow and to fix N2.


Photosynthetic cyanobacteria are also N-fixing organisms and are especially important in paddy rice (Oryza). The amount of N fixed varies greatly from crop to crop, ranging from a few kg to a few hundred kg N ha1 yr1. The process is depressed by ample N supply from other sources, and it is sensitive to lack of phosphorus. The amount of globally fixed N is almost the double the amount of applied fertilizer N. Next to symbiotic N fixing bacteria also non-symbiotic species (e.g. Azotobacter) occur in soils. In general, free-living diazotrophs make a small but significant contribution to the soil N status. Some nonleguminous trees and plants (e.g. alder (Alnus), sugarcane (Saccharum) host N-fixing bacteria as well. Much uncertainty exists about the association of N fixing bacteria with non-legumes (so called associative N fixing bacteria). Mineral and Organic Nitrogen Fertilization Theoretically plants should prefer NH4þ above NO3, because NH4þ does not need to be reduced before incorporation into the plant. In most well-drained soils oxidation of NH4þ is fairly rapid and therefore most plants have developed to grow better with NO3. However, a number of studies have shown that plants better develop when both sources are available. Rice, growing under submerged conditions must grow in the presence of NH4þ as NO3 is not stable under flooded conditions. When urea is applied it rapidly hydrolyzes under well-drained conditions, unless a urease inhibitor is being added; under submerged conditions rice plants may also absorb N directly as molecular urea. Organic manure can be of plant or animal origin or a mixture of both. However, most comes from dung and urine from farm animals. It exists as farmyard or stable manure, urine, slurry or as compost. Because its composition is not constant and because plant material (catch or cover crops, legumes) is often added freshly (green manure) to the soil, less than 30% of its nutrients becomes available for the next crop.

NITROGEN UPTAKE BY PLANTS Growing plants get their N from fertilizer N as well as from organic soil N upon mineralization. Plants take up N compounds both as NO3 and as NH4þ. In general, NO3 is the major source of plant N. There is some evidence that small amounts of organic N (urea or amino acids) can be taken up by plants from the soils solution. Plant uptake of N can be studied through the use of mineral fertilizers or organic matter labeled with the stable N isotope 15N. The proportion of applied N taken up by the crop is affected by


1127

many factors, including crop species, climate and soil conditions. Above ground parts of the crop can recover 40–60% of the fertilizer N applied.

NITROGEN LOSS PROCESSES Ammonia Volatilization Losses of N from the soil by NH3 volatilization amount globally to 54 Mt (or 1012 g) NH3-N yr1 and 75% is of anthropogenic origin.[4] According to the ECETOC,[5] the dominant source is animal manure and about 30% of N in urine and dung is lost as NH3. The other major source is surface application of urea or ammonium bicarbonate and to a lesser degree other ammonium-containing fertilizers. As urea is the most important N fertilizer in the world, it may lead to important NH3 loss upon hydrolysis and subsequent pH rise in the vicinity of the urea till. The transformation of NH4þ to the volatile form NH3 increases with increasing pH, temperature, soil porosity, and wind speed at the soil surface. It decreases with increasing water content and rainfall events following application. Ammonia losses from soils can be effectively reduced by fertilizer incorporation or injection instead of surface application. Emission of Nitrogen Oxides (N2O, NO) and Molecular Nitrogen (Nitrification and Denitrification) Microbial nitrification and denitrification are responsible for the emission of NO and N2O.[6] They are by-products in nitrification and intermediates during denitrification. Probably about 0.5% of fertilizer N applied is emitted as NO[7] and 1.25% as N2O.[8] However, wide ranges have been reported. Intensification of arable agriculture and of animal husbandry has made more N available in the soil N cycle increasing the emission of N oxides. The relative percentage of NO and N2O formation very much depends on the moisture content of the soil. At a water-filled pore space (WFPS, or the fraction of total soil pore space filled with water) below 40% NO is produced mainly from nitrification. Between a WFPS of 40% and 60% formation of NO and N2O from nitrification occurs. Between a WFPS of 60% and 80% N2O is predominantly produced from denitrification and the formation of NO is decreasing sharply. At a WFPS above 80% the formation of N2 by denitrification is dominant. In practice these WFPS ranges will overlap anddepend on the soil type.[9] Next to water content, also temperature, land use and availability of N and decomposable organic matter are important determining factors for N2O formation. Nitrous oxide is a greenhouse

1128

gas contributing 5–6% to the enhanced greenhouse effect. Increased concentrations are also detrimental for the stratospheric ozone layer.[10] In the presence of sunlight, NOx (NO and NO2) react with volatile organic compounds from evaporated petrol and solvents and from vegetation and forms tropospheric ozone which is, even at low concentration, harmful to plants and human beings. The major gaseous end-product of denitrification is N2. The ratio of N2O to N2 produced by denitrification depends on many environmental conditions. Generally themore anaerobic the environment the greater the N2 production. Denitrification is controlled by three primary factors (oxygen, nitrate and carbon), which in turn are controlled by several physical and biological factors. Denitrification N loss can reach 10% of the fertilizer N input—more on grassland and when manure is also applied.[11] Chemical denitrification is normally insignificant and is mainly related to the stability of NO2 and acid conditions.[12] It is more difficult to reduce N2O and NO from soils then NH3 losses. A general principle is to minimize N surpluses in the soil profile via carefulfertilizer adjustment, corresponding to the actual crop demands. Leaching Applied NO3 or NO3, formed through nitrification from mineralized NH4þ or from NH4þ from animal manure, can leach out of the rooting zone. It is well possible that this leached NO3 can be denitrified at other places and returned into the atmosphere. The amount and intensity of rainfall, quantity and frequency of irrigation, evaporation rate, temperature, soil texture and structure, type of land use, cropping and tillage practices and the amount and form of fertilizer N are all parameters influencing the amount of NO3 leaching to the underground water. Nitrate leaching should be kept under control as it may influence the nitrate content in drinking water influencing human health and in surface water, causing eutrophication. Nitrate losses can be minimized by reducing the mineral N content in the soil profile during the



winter period by careful fertilizer adjustment, growing of cover crops or riparian buffer areas.

REFERENCES 1. McCarty, G.W. Modes of action of nitrification inhibitors. Biol. Fert. Soils 1999, 29, 1–9. 2. Prosser, J.I., Ed. Nitrification, Special Publications of the Society of General Microbiology; IRL Press: Oxford, 1986; 20 pp. 3. Lagreid, M.; Bockman, O.C.; Kaarstad, O. Agriculture, Fertilizers and the Environment; CIBA Publishing: Oxon, UK, 1999; 294 pp. 4. Sutton, M.A.; Lee, D.S.; Dollard, G.J.; Fowler, D. International conference on atmospheric ammonia: emission, deposition and environmental impacts. Atmospheric Environment 1998, 32, 1–593. 5. ECETOC Ammonia Emissions to Air in Western Europe, (No. 62); European Centre for Ecotoxicology and Toxicology of Chemicals: Brussels, 1994; 196 pp. 6. Bremner, J.M. Sources of nitrous oxide in soils. Nutr. Cycl. Agroecosys. 1997, 49, 7–16. 7. Veldkamp, E.; Keller, M. Fertilizer-induced nitric oxide emissions from agricultural soils. Nutr. Cycl. Agroecosys. 1997, 48, 69–77. 8. Mosier, A.; Kroeze, C.; Nevison, C.; Oenema, O.; Seitsinger, S.; Van Cleemput, O. Closing the global N2O budget: nitrous oxide emissions through the agricultural N cycle. Nutr. Cycl. Agroecosys. 1998, 52, 225–248. 9. Davidson, E.A. Fluxes of Nitrous Oxide and Nitric Oxide from terrestrial ecosystems. In Microbial Production and Consumption of Greenhouse Gases: Methane, Nitrogen Oxides, and Halomethanes; Rogers, J.E., Whitman, W.B., Eds.; American Society for Microbiology: Washington, DC, 1991; 219–235. 10. Crutzen, P.J. The influence of nitrogen oxides on the atmospheric ozone content. Quat. J. Royal Meteor. Soc. 1976, 96, 320–325. 11. von Rheinbaben, W. Nitrogen losses from agricultural soils through denitrification—A critical evaluation. Z. Pflanzenern. Bodenk. 1990, 153, 157–166. 12. Van Cleemput, O. Subsoils: chemo- and biological denitrification, N2O and N2 emissions. Nutr. Cycl. Agroecosys. 1998, 52, 187–194.

Nitrous Oxide Emissions: Agricultural Soils John R. Freney Commonwealth Scientific and Industrial Research Organisation (CSIRO) Plant Industry, Canberra, Australian Capital Territory, Australia

INTRODUCTION Nitrous oxide is a gas that is produced naturally by many different micro-organisms in soils and waters, and as a result of human activity associated with agriculture, biomass burning, stationary combustion, automobiles, and the production of nitric and adipic acids for industrial purposes. According to the Intergovernmental Panel on Climate Change (IPCC),[1] 23.1 million metric tons (Mt) of nitrous oxide is emitted each year, 14.1 Mt as a result of natural processes (4.7 Mt from the oceans, 6.3 Mt from tropical soils, and 3.1 Mt from temperate soils); and 9 Mt as a result of human activities (5.5 Mt from agricultural soils, 0.6 Mt from cattle and feedlots, 0.8 Mt from biomass burning, and 2.1 Mt from mobile sources and industry). While there is considerable uncertainty associated with each of these estimates, it is apparent that most nitrous oxide is derived from soils. Because of the intimate connection between the Earth and the atmosphere, much of the nitrous oxide produced enters the atmosphere and affects its chemical and physical properties. Nitrous oxide contributes to the destruction of the stratospheric ozone layer that protects the Earth from harmful ultraviolet radiation, and is one of the more potent greenhouse gases that trap part of the thermal radiation from the Earth’s surface. The atmospheric concentration of nitrous oxide is 313 parts per billion. It is increasing at the rate of 0.7 parts per billion each year, and its lifetime is 166 years.[2] It seems that the increased atmospheric concentration results from the increased use of synthetic fertilizer nitrogen, biologically fixed nitrogen, animal manure, crop residues, and human sewage sludge in agriculture to produce food and fiber for the rapidly increasing world population.[3]

NITROUS OXIDE EMISSION FROM AGRICULTURE All soils are deficient in nitrogen for the growth of plants, but the deficiency can be overcome by adding fertilizer nitrogen. When the fertilizer (e.g., urea or Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001564 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

ammonia-based compounds) is applied to soil, it is transformed by micro-organisms as follows: 1

2

Fertilizer nitrogen ! Ammonium ! Nitrite 3

4

! Nitrate ! Nitrite 5

6

! Nitric oxide ! Nitrous oxide 7

! Dinitrogen

ð1Þ

When the soil is aerobic (i.e., when oxygen is present) ammonium is oxidized to nitrite and nitrate (Steps 2 & 3). This process is called nitrification. After addition of irrigation water or rain, the soil may become anaerobic (devoid of oxygen). The nitrate is then reduced by soil organisms to nitrite and the gases nitric oxide, nitrous oxide, and dinitrogen (Steps 4–7) in a process termed denitrification.[4] When atmospheric scientists first expressed concern that nitrous oxide emission into the atmosphere, as a result of fertilizer use, would lead to destruction of the ozone layer, it was thought that nitrous oxide was produced mainly from the microbiological reduction of nitrate in poorly aerated soils. However, research in the latter part of the 1970s showed that significant nitrous oxide was emitted from aerobic soils during nitrification of ammonium, and subsequent work has shown that nitrification is a major source of nitrous oxide.[4]

Nitrous Oxide from Denitrification Certain micro-organisms in the absence of oxygen have the capacity to reduce nitrate (or other nitrogen oxides). Most denitrifying bacteria are heterotrophs—that is, they require a source of organic matter for energy— but denitrifying organisms that obtain their energy from light or inorganic compounds also occur in soils. The capacity to denitrify has been reported in more than 20 genera of bacteria, and almost all are aerobic organisms that can only grow anaerobically in the presence of nitrogen oxides. The dominant denitrifying organisms in soil are Pseudomonas and Alcaligenes. In addition to the free-living denitrifiers, Rhizobia 1129

1130

living symbiotically in root nodules of legumes are able to denitrify nitrate and produce nitrous oxide.[4] The general requirements for biological denitrification include the presence of micro-organisms with denitrifying capacity, nitrate (or other nitrogen oxides) and available organic matter, the absence of oxygen, and a suitable pH and temperature environment. In aerobic soils, denitrification can occur in anaerobic microsites in soil aggregates or in areas of high carbon content, where active microbial activity rapidly consumes all of the available oxygen.[4] Nitrous Oxide from Nitrification The process of nitrification is normally defined as the biological oxidation of ammonium to nitrate with nitrite as an intermediate.[4] The first step in the reaction, the oxidation of ammonium to nitrite, is carried out mainly by the micro-organism Nitrosomonas. The second step, oxidation of nitrite to nitrate, is carried out by Nitrobacter. It has been shown in a number of publications that Nitrosomonas europaea produces nitrous oxide during the oxidation of ammonium.[4] The possibility that significant nitrous oxide can be produced in soils by nitrifying organisms was indicated by studies that showed that soils incubated under aerobic conditions with ammonium produced more nitrous oxide than soils amended with nitrate.[4] In addition, treatment of aerobic soils with nitrapyrin, which inhibits nitrification of ammonium but has little effect on denitrification, markedly reduced the emission of nitrous oxide.[4] Production of nitrous oxide by nitrification in soils is increased by increasing temperature, pH, water content, available carbon, and the addition of ammonium-based fertilizers, plant residues, and animal manure. Flooded Soils In the past few years, increased attention has been given to nitrous oxide emission from paddy soils. The concern is that the introduction of management practices to reduce methane emissions from flooded soils may result in increased emissions of nitrous oxide. Flooded soils are characterized by an oxygenated water column overlying an oxidized layer at the soil– water interface, an aerobic zone around each root, and anaerobic conditions in the remainder of the soil. This differentiation of the flooded soil into oxidized and reduced zones has a marked effect on the transformation of nitrogen.[5] The resulting reactions are as follows: 1. Ammonium in the reduced zone diffuses to the oxidized zone;


Nitrous Oxide Emissions: Agricultural Soils

2. Ammonium is oxidized to nitrate by nitrifying organisms; 3. The nitrate formed diffuses to the anaerobic zone; 4. Denitrification occurs with the production of nitrous oxide and dinitrogen; 5. The gaseous products diffuse through the soil and water layers to the atmosphere.[6] It is apparent that the rate of diffusion of ammonium to an oxidized layer and the rate of nitrification in the oxidized layer are factors controlling the production of nitrous oxide in flooded soils. The rate of diffusion of nitrous oxide through the soil and water layers will control its rate of emission to the atmosphere, or its further reduction to dinitrogen.[5] A number of mechanisms have been identified for the transfer of nitrous oxide from the soil to the atmosphere.[3] Nitrous oxide may diffuse from the zone of production through the saturated soil and water layer to the atmosphere. It may also enter the roots of the rice plant and move by diffusion through the plant to the atmosphere in the same way as methane. Bhadrachalam et al.[6] studied the importance of the two pathways in intermittently flooded rice in the field in India and found that nitrogen gas fluxes were 30% greater when transfer through the plants was included. In the tropics, rice is usually transplanted and fertilized some time after flooding. Because of the anaerobic conditions that develop before fertilization and the slow rate of diffusion of nitrous oxide in flooded soils, most of the nitrous oxide is reduced to dinitrogen and very little escapes to the atmosphere. Nitrous oxide emission from intermittently flooded rice was relatively large compared with that from permanently flooded rice, reflecting the different oxidation states of intermittently and continuously flooded soils.[6] Studies of nitrous oxide emission from rice fields from the time the soils were drained for harvest, through to flooding the soil in preparation for planting the next crop, showed that nitrous oxide was emitted continuously while the soil was not flooded. Overall, the rate of emission of nitrous oxide from floodedsoils was less than that from upland soils after application of nitrogen fertilizer.[3]

Biomass Burning During combustion the nitrogen in the fuel can be converted into gaseous forms such as ammonia, nitric oxide, nitrous oxide, dinitrogen, and hydrogen cyanide. It is estimated that biomass burning contributes between 0.3 and 1.6 Mt nitrous oxide per year globally to the atmosphere.[3] Most of the biomass burning


1131

(90%) takes place in the tropics as a result of forest clearing, savanna and sugar cane fires, and burning of agricultural wastes and firewood.[7] Biomass burning is not only an instantaneous source of nitrous oxide, but it results in a longer-term enhancement of the production of this gas. Measurements of nitrous oxide emissions from soils, before and after burning showed that significantly more nitrous oxide was exhaled after the burn through alteration of the chemical, biological, and physical processes in soil.[7]

Table 1 Calculated emission of nitrous oxide from agricultural activities Mt nitrous oxide per year Direct soil emissions Synthetic fertilizer Animal waste Biological nitrogen fixation Crop residue Cultivation of Histosols Total

1.4 (0.28–2.5) 0.9 (0.19–1.7) 0.16 (0.03–0.3) 0.6 (0.11–1.1) 0.16 (0.03–0.3) 3.3 (0.6–5.9)

Animal production Waste management systems

3.3 (0.9–4.9)

Indirect emissions

Fertilizer Consumption and Nitrous Oxide Production Nitrous oxide emissions from agricultural soils are generally greater and more variable than those from uncultivated land. Application of fertilizer nitrogen, animal manure, and sewage sludge usually results in enhanced emissions of nitrous oxide.[7] Generally, there is a large emission of nitrous oxide immediately after the application of fertilizer. After about 6 weeks, the emission rate falls and fluctuates around a low value. Mosier[8] concluded that interactions between the physical, chemical, and biological variables are complex, that nitrous oxide fluxes are variable in time and space, and that soil management, cropping systems, and variable rainfall appear to have a greater effect on nitrous oxide emission than the type of nitrogen fertilizer. Consequently, Mosier et al.[9] recommend the use of one factor only for calculating the emission of nitrous oxide from different fertilizer types: N2 O emitted ¼ 1:25% of N applied ðkg N=haÞ

ð2Þ

This equation is based on data from long-term experi-ments with a variety of mineral and organic fertilizers, and encompasses 90% of the direct contributions of nitrogen fertilizers to nitrous oxide emissions. Mosier et al.[3] developed a methodology to estimate agricultural emissions of nitrous oxide, taking into account all of the nitrogen inputs into crop production. They included direct emissions from agricultural soils as a result of synthetic fertilizer addition, animal wastes, increased biological nitrogen fixation, cultivation of mineral and organic soils through enhanced organic matter mineralization, and mineralization of crop residues returned to the field. Indirect nitrous oxide emissions resulting from deposition of ammonia and oxides of nitrogen, leaching of nitrate, and introduction of nitrogen into sewage systems were also included. They concluded that in 1989, 9.9 Mt of


Atmospheric deposition Nitrogen leaching and runoff Human sewage Total Total

0.47 (0.09–0.9) 2.5 (0.2–12.1) 0.3 (0.06–4.1) 3.3 (0.35–17.1) 9.9 (1.9–27.9)

(Modified from Ref.[3].)

nitrous oxide was emitted into the atmosphere directly or indirectly, as a result of agriculture (Table 1). MANAGEMENT PRACTICES TO DECREASE NITROUS OXIDE EMISSION The low efficiency of fertilizer nitrogen in agricultural systems is primarily caused by the large losses of mineral nitrogen from those systems by gaseous loss: nitrous oxide emission is directly linked to the loss processes. It is axiomatic that any strategy that increases the efficiency of nitrogen fertilizer use will reduce emissions of nitrous oxide, and this has been directly demonstrated for a number of strategies.[3] The IPCC[1] reported that some combination of the following management practices, if adopted worldwide, would improve the efficiency of the use of synthetic fertilizer and manure nitrogen, and significantly reduce nitrous oxide emission into the atmosphere: 1. Match nitrogen supply with crop demand. 2. Tighten nitrogen flow cycles by returning animal wastes to the field and conserving residues instead of burning them. 3. Use controlled-release fertilizers, incorporate fertilizer to reduce volatilization, use urease and nitrification inhibitors, and match fertilizer type to precipitation. 4. Optimize tillage, irrigation, and drainage. The potential decrease in nitrous oxide emissions from synthetic fertilizer, as a result of the mitigation techniques, could amount to 20%.[1]

1132


REFERENCES 1. IPCC (Intergovernmental Panel on Climate Change). Climate Change 1995. Impacts, Adaptations and Mitigation of Climate Change: Scientific–Technical Analyses; Watson, R.T., Zinyowera, M.C., Moss, R.H., Eds.; Cambridge University Press: Cambridge, England, 1996; 1–878. 2. Hengeveld, H.; Edwards, P. 1998. An assessment of new research developments relevant to the science of climate change. Climate Change Newsletter 2000, 12, 1–52. 3. Mosier, A.; Kroeze, C.; Nevison, C.; Oenema, O.; Seitzinger, S.; van Cleemput, O. Closing the global N2O budget: nitrous oxide emissions through the agricultural nitrogen cycle. Nutri. Cycling Agroecosys. 1998, 52, 225–248. 4. Bremner, J.M. Sources of nitrous oxide in soils. Nutri. Cycling Agroecosys. 1997, 49, 7–16. 5. Patrick, W.H., Jr. Nitrogen transformation in submerged soils. In Nitrogen in Agricultural Soils; Stevenson, F.J.,


6.

7.

8.

9.

Ed.; American Society of Agronomy: Madison, WI, 1982; 449–465. Bhadrachalam, A.; Chakravorti, S.P.; Banerjee, N.K.; Mohanty, S.K.; Mosier, A.R. Denitrification in intermittently flooded rice fields and N-gas transport through rice plants. Ecol. Bull. 1992, 42, 183–187. Granli, T.; Bøckman, O.C. Nitrous oxide from agriculture. Norwegian J. Agric. Sci. 1994, Supplement No. 12, 1–128. Mosier, A.R. Chamber and isotope techniques. In Exchange of Trace Gas between Terrestrial Ecosystems and the Atmosphere; Andreae, M.O., Schimel, D.S., Eds.; John Wiley & Sons: Chichester, England, 1989; 175–187. Mosier, A.R.; Duxbury, J.M.; Freney, J.R.; Heinemeyer, O.; Minami, K. Nitrous oxide emissions from agricultural fields: assessment, measurement and mitigation. Plant Soil 1995, 181, 95–108.

Nitrous Oxide Emissions: Sources, Sinks, and Strategies Katsu Minami National Institute for Agro-Environmental Sciences, Tsukuba, Japan

INTRODUCTION

Sinks for Nitrous Oxide

Although it has been known for more than 50 years that nitrous oxide (N2O) is a regular constituent of the atmosphere, it was not considered to be of any importance as an air constituent until the early 1970s. Atmospheric scientists hypothesized that N2O released to the atmosphere through denitrification of nitrates in soils and natural waters may trigger reactions in the stratosphere leading to partial destruction of ozone layer protecting the earth from biologically harmful ultraviolet radiation from the sun.[1] Nitrous oxide is also one of the natural components of Earth’s atmosphere and contributes to the natural greenhouse effect; therefore the increasing of N2O in the atmosphere may be contributing to global warming.[2] The atmospheric concentration of N2O has been increasing at an accelerated rate for several decades at a rate of 0.8 ppbv=yr, and the lifetime of N2O is 120 yr. It has been estimated that doubling the concentration of N2O in the atmosphere would result in a 10% decrease in the ozone layer which would increase the ultraviolet radiation reaching the earth by 20%,[3] eventually leading to an increase in the occurrence of skin cancer and other health problems. The global warming potential (GWP) of each molecule of N2O is about 210 times (20-year horizon) greater than each molecule of CO2. Nitrous oxide currently accounts for 6% of total GWP.[4]

The major atmospheric loss process for N2O is photochemical decomposition by sunlight (wavelengths 180–230 nm) in the stratosphere and is calculated to be 12.3 (range 9–16) Tg N=yr. Tropospheric sinks such as surface loss in aquatic and soil systems are considered to be small as compared to atmospheric sink.[5] However, the paucity of data does not enable us to estimate the importance of this sink on a global basis.

THE BUDGET OF ATMOSPHERIC NITROUS OXIDE

Sources of Nitrous Oxide There are many sources of nitrous oxide, both natural and anthropogenic, which cannot be easily quantified. Undoubtedly, the earth and oceans are significant N2O sources. Nitrous oxide fluxes from upwelling regions of the Indian and Pacific Oceans clearly suggest that oceans may be a larger source. The ocean flux estimate is 3 (range 1–5) Tg N=yr. Tropical forest soils are probably the single most important source of N2O emission to the atmosphere. The total N2O emission from tropical soils (forest, savannah) is estimated at 4 (range 2.7–5.7) Tg N=yr. The magnitude of N2O emissions from intensively fertilized tropical agricultural soils has not been quantified. Also, no attempt has been made to separate the tropical soil sources into natural and anthropogenic components.[5] The main anthropogenic sources are derived from agriculture and a number of industrial processes such as adipic acid and nitric acid production. New research suggests that N2O emissions from cropped, nitrogenfertilized agricultural systems are significant on a global scale as shown in Table 1.

Atmospheric Distribution of Nitrous Oxide As a result of biotic and anthropogenic activities, the concentration of N2O in the atmosphere is increasing at the rate of about 0.25%=yr. The concentration of N2O is about 0.75 ppbv higher in the northern hemisphere than in the southern hemisphere, suggesting the presence of greater source of N2O in the northern hemisphere than in the southern hemisphere. Ice core measurements show that the preindustrial value of N2O was relatively stable at about 285 ppbv for most of the past 2000 years, and started to increase around the year 1700 associated with anthropogenic activity.[4,5] Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042718 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

Agricultural Fields About 40% of N2O sources are anthropogenic and, among them, fertilized soils account for about 60%.[5] This figure could be an underestimate because tropical agricultural soil sources resulting from human activities have not been separated from natural tropical soil sources. Agricultural N2O emissions are considered to arise from fertilization of soils with mineral N and animal manures, N derived from biological N fixation, and from enhanced soil N mineralization. Nitrous oxide is 1133

1134

Nitrous Oxide Emissions: Sources, Sinks, and Strategies

Table 1 The amount of N2O emission from agriculture (Mt=yr) Mineral fertilizer N-fixation Soils after burning Animal wastes Biomass burning Forest conversion

1.5 0.5 0.1 1.5 0.2 0.4

(0.5–2.5) (0.25–0.75) (0.05–0.2) (0.5–2.5) (0.1–0.3) (0.1–1)

Total

4.2 (1.5–7.25)

oxygen concentrations. Because oxygen supply is moderated by soil moisture, the effect of soil water content on N transformation probably reflects its impact on oxygen diffusion in the normal soil moisture range. Nitrous oxide emissions from N-fertilizer applied agricultural fields have been detected by Bremner and Blackmer.[7]

Denitrification directly evolved during biomass burning, and produced in soil after burning, and enhanced emissions arise during the clearing of tropical forests for agricultural activities.[4] Bouwman[6] estimated the total emissions of N2O from a regression equation: total annual direct field N2O–N Loss ¼ 1 þ 0.0125 N-application (Kg N=ha). The value of 1 Kg N2O–N=ha represents the background N2O-N evolved and the 0.0125 factor expresses for the contribution from fertilization. This estimate includes N sources from a variety of mineral and organic N fertilizers and was based on long-term data sets. About 40% of the estimated N2O production is derived from North and Central America, Europe, and the former Soviet Union where about 20% of the world human population resides. Asian countries that hold about 55% of the global human population contribute about 40% of the estimated annual N2O production.[4]

MECHANISM OF NITROUS OXIDE PRODUCTION

Biological denitrification refers to the dissimilatory reduction of nitrate and nitrite to produce NO, N2O, and N2 by a taxonomically diverse group of bacteria which synthesize a series of reductases that enable them to utilize successively more reduced N oxides as electron acceptors in the absence of oxygen. The general reductive sequence is as follows NO3 ! NO2 ! NO ! N2 O ! N2 The most abundant denitrifiers are heterotrophs that require sources of electron-reducing equivalents contained in available organic matter. Soil factors that most strongly influence denitrification are oxygen, nitrate concentration, pH, temperature, and organic carbon. Nitrous oxide reductase appears to be more sensitive to oxygen than either nitrate or nitrite reductase. Therefore, N2 production predominates in more anoxic sites and N2O production may be higher under more aerobic conditions.

Chemical Decomposition of Nitrite (Chemodenitrification)

Nitrification Nitrification is the reaction whereby ammonium is oxidized to nitrate. In soils, autotrophic and heterotrophic bacteria mediate this process. The most common ammonium oxidizers are Nitrosomonas spp., which are involved in the formation of nitrite, while nitrite oxidation to nitrate is usually achieved by Nitrobacter spp. The overall nitrification sequence is as follows N2 O " NHþ 4 ! NH2 OH ! ðNOHÞ

! NO ! NO2 ! NO3

Two nitrogenous gases may evolve through nitrification, NO and N2O. The nitrifiers are active over a wide range of temperatures (2–40 C). The overall nitrification process is controlled primarily by ammonium and


There is evidence that nitrite produced by nitrifying or denitrifying micro-organisms can also react chemically to form N2O via ‘‘chemodenitrification.’’[8] High nitrite concentrations have been attributed to the inhibition of nitrite oxidation, which is presumed to result from ammonia toxicity to Nitrobacter. Several investigators have noted that gaseous loss of nitrogen (via NO, N2O, or N2) may accompany temporary nitrite accumulation. High concentration of nitrite is sometimes found in anaerobic soils where ammonium and ammonium type fertilizers are applied at high doses. Nitrite ions react chemically with organic molecules forming nitroso-groups (–N¼O) that are unstable.

Other Mechanisms Some of the N2O evolved from soils may be formed by chemical decomposition of hydroxylamine (NH2OH)


produced by nitrifying or nitrate-reducing microorganisms, because NH2OH has been identified as an intermediate in oxidation of ammonium to nitrate by Nitrosmonas europeae and has been postulated as an intermediate in microbial reduction of nitrate to ammonium, and several investigations have shown that NH2OH is decomposed rapidly in soils with formation of N2O and N2 by processes that appear to be largely chemical. Other investigations indicated that Mn compounds are involved in the reactions leading to formation of N2O and N2 by chemical decomposition of NH2OH in soils, and that CaCO3 and Fe compounds are involved in the reactions leading to formation of N2 in calcareous soils. The reaction is as follows 2MnO2 þ 2NH2 OH ! 2MnO þ N2 O þ 3H2 O Several workers[1] have postulated that N2O is produced in soils through interaction of hydroxylamine and nitrite produced by soil micro-organisms as follows NH2 OH þ HNO2 ! N2 O þ 2H2 O

1135

Nitrification Inhibitors Because ammonia and ammonium producing compounds are the main sources of fertilizer nitrogen, maintenance of the applied nitrogen in the ammonium form should result in a lower emission of N2O from cultivated soils. One mechanism of maintaining added ammonium N is to use nitrification inhibitors with the fertilizer. Mosier et al.[12] summarized the effects of nitrification inhibitors on N2O emission from fertilized soils both in laboratory and field studies. A number of field studies indicate that nitrification inhibitors do limit N2O emissions from ammonium-based fertilizers. Several recent field tests also show that the utilization of a variety of nitrification inhibitors does significantly limit N2O emissions from the application of ammonium-based fertilizers. To illustrate this point, a study to quantify the effect of nitrification inhibitors DCS (N-2, 5-dichlorophenyl succinamic acid) and the application of ammonium sulfate on N2O emissions was conducted in field lysimeters using carrot (Daucus carota L.) as a test crop.[11] The addition of DCS reduced about 30% of N2O emission and leaching of nitrate.

Controlled Release Fertilizer MEASUREMENT OF NITROUS OXIDE EMISSION Nitrous oxide emissions from N-fertilized agricultural fields have been found to vary between 0.001% and 6.8% of the N applied to the field.[9–11] A proportion of this variability in N2O estimates, relative to the amount and type of fertilizer applied and to the type of cropping system such as grassland, upland crop, rice paddy, and others, has been attributed to spatial and temporal change in the processes, which produce N2O in soil.

MITIGATION STRATEGY Many of the strategies were proposed by Mosier et al.[12] 1) match N supply with crop demand; 2) close N flow cycles; 3) use advanced fertilizer technologies; and 4) optimize tillage, irrigation, and drainage, in which gaseous emissions deal primarily with cropping systems could be minimized. Although most of the practices listed are assumed to decrease N2O emissions, there have been relatively few systematic studies in which various farming practices were compared as to their ability to conserve N and limit N2O emissions. A number of field studies have been conducted with nitrification inhibitors that could decrease N2O emissions when used. There are a few studies in which the potential of using controlled release fertilizer for decreasing N2O emission was evaluated as follows.


The use of controlled release fertilizer has the potential to improve N-use efficiency by matching nutrient release with crop demand, reducing NO3 leaching and denitrification losses. Polyolefin-coated fertilizers are a type of controlled release fertilizer where fertilizer granules are covered with a thermoplastic resin. The release of the N fertilizer is temperature dependent and is not controlled by hydraulic reactions or by microbial attack of the coating. Greenhouse studies have revealed that controlled release fertilizer can increase yields with more N-fertilizer use efficiency and reduce NO3 leaching. For example, Minami[11] observed that a controlled release fertilizer reduced N2O emissions in lysimeter studies of carrot at Tsukuba, Japan, in which the fertilizer-induced emissions of N2O–N during an 83 day-period of cultivation were 0.14% and 0.02% of the 250 Kg N applied following ammonium sulfate and a controlled release fertilizer, respectively.

REFERENCES 1. Bremner, J.M. Sources of nitrous oxide in soils. Nutr. Cycling in Agroecosys. 1997, 49, 7–16. 2. IPCC. Climate Change: The IPCC Scientific Assessment; Houghton, J.T., Jenkins, G.J., Ephraume, J.J., Eds.; Cambridge University Press: Cambridge, 1990.

1136

3. Crutzen, P.J.; Ehhalt, D.H. Effect of nitrogen fertilizers and combustion on the stratospheric ozone layer. Ambio 1977, 6, 112–117. 4. IPCC. In Climate Change 1995, Impacts, Adaptations and Mitigation of Climate Change: Scientific-Technical Analyses; Watson, R.T., Zinyowera, Mos, R.H., Eds.; Cambridge University Press: Cambridge, 1996. 5. IPCC. In Climate Change 1994: Radiative Forcing of Climate Change and an Evaluation of the IPCC IS92 Emission Scenarios; Houghton, J.T., Meira Filho, L.G., Bruce, J., Lee, H., Callander, B.A., Haites, E., Harris, N., Maskell, K., Eds.; Cambridge University Press: Cambridge, 1995. 6. Bouwman, A.F. Direct Emission of Nitrous Oxide from Agricultural Soils, RIVM Report No. 773004004; RIVM: Bilthoven, 1994; 1–28. 7. Bremner, J.M.; Blackmer, A.M. Effects of acetylene and soil water content on emissions of nitrous oxide from soils. Nature 1979, 280, 380–381.



8. Chalk, P.K.; Smith, C.J. Chemodenitrification. In Gaseous Loss of Nitrogen from Plant-Soil Systems; Freney, J.R., Simpson, J.R., Eds.; Martinus Nijhof=Dr W. Junk Publishers: The Hague, 1983; 65–90. 9. Eichner, M.J. Nitrous oxide emissions from fertilized soils: summary of available data. J. Environ. Qual. 1990, 19, 272–280. 10. Bouwman, A.F. Exchange of greenhouse gases between terrestrial ecosystems and the atmosphere. In Soils and the Greenhouse Effect; Bouwman, A.F., Ed.; John Wiley & Sons: New York, 1990; 61–127. 11. Minami, K. Nitrous oxide emissions from agricultural fields. In Trace Gas Emissions and Plants; Singh, S.N., Ed.; Kluwer Academic Publishers: Dordrecht, 2000; 215–230. 12. Mosier, A.R.; Duxbury, J.M.; Freney, J.R.; Heinemeyer, O.; Minami, K. Nitrous oxide emissions from agricultural fields: assessment, measurement and mitigation. Plant Soil 1996, 131, 95–108.

No Till Paul W. Unger United States Department of Agriculture (USDA), Bushland, Texas, U.S.A.

INTRODUCTION No-till, also known as no-tillage or zero-tillage, is a type of conservation tillage, which is a planting system that results in at least 30% cover of crop residues on the soil surface after planting the next crop.[1] Use of conservation tillage provides major soil erosion control benefits and also helps conserve water. In contrast, conventional tillage refers to tillage operations normally used for crop production that bury most residues and result in proteins > cellulose > hemicelluloses > fats, starches, and waxes > lignins and tannins. Mineralization releases soluble or gaseous inorganic constituents during decomposition processes. Humification is a multistage process.[3] Source materials for humus synthesis include residual components from incomplete decomposition of organic litter and the products of microbial anabolic activities. According to present concepts, polyphenols derived from lignin degradation, together with those synthesized by microorganisms, are oxidized to quinones, which undergo self-polymerization or combine with amino compounds to form nitrogen (N)-containing polymers. Sugar-amine condensation reactions may also participate in the formation of humic substances. Accumulation in Organic Soils Paludization can be considered a geogenic rather than pedogenic process because it involves deposition of initial parent material. Paludization occurs when conditions impede decomposition, enabling the buildup of a thick mass of organic deposits. Decomposition is hampered by poorly drained conditions, as in Histosols, and by extreme cold, as in Gelisols. Under anaerobic conditions, humic substances exhibit an accumulation of aromatic carbon (C) compounds arising from the absence of lignin-degrading fungi.[4] Aromaticity can also develop in organic horizons from sources without lignin, such as detritus from algae and mosses. In contrast to paludization, ripening refers to the decomposition processes occurring in the organic horizon under oxidizing conditions after exposure to the air. Surface Accumulation in Mineral Soils

Organic Matter Transformation Decomposition refers to the chemical and biochemical reactions occurring during the decay of plant and animal remains as soil microorganisms colonize them (Fig. 1). Decomposition involves the fragmentation 1172 Copyright © 2006 by Taylor & Francis

Melanization produces thick, dark-colored surface horizons characteristic of Mollisols. The formation of mollic epipedons is promoted by the proliferation of grass roots that constitute a considerable input of plant residues.[5] Another key factor in melanization Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001950 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter Accumulation

1173

Table 1 Fundamental pedogenic processes associated with organic matter accumulation Term

Definition

Representative horizon (soil order)

Littering

Accumulation of fresh organic detritus on the mineral soil surface to a depth of 30 cm deep) organic materials on the mineral soil surface

Histic epipedon (Histosols and Gelisols)

Ripening

Changes in organic soil promoted by entry of air into previously waterlogged material

Histic epipedon (Histosols)

Melanization

Darkening of light-colored initial mineral soil by addition of organic matter

Mollic epipedon (Mollisols) Melanic epipedon (Andisols)

Podzolization

Translocation of organic matter in the soil profile associated with Al and Fe migration

Spodic horizon (Spodosols)

(From Ref.[2].)

is the active faunal community, which contributes to the rapid incorporation of the residues into the mineral soil and favors high initial mineralization rates. Subsequent decomposition and humification processes result in the formation of chemically stable, darkcolored humic substances, characterized by a high proportion of high-molecular weight, highly aromatic, acid-insoluble humic acids. Multivalent cations, such as calcium, act as bridges between organic colloids and clay particles, and stabilize organic substances within the soil matrix. In allophanic soils derived from volcanic parent material (Andisols), organic matter accumulation is favored by the formation of resistant organic-alumina complexes.

Subsurface Accumulation Podzolization results in the formation of subsurface horizons of organic matter accumulation characteristic of Spodosols. Subsurface organic matter accumulation occurs at the top of the spodic horizon due to the migration of water-soluble organic compounds from the mineral surface.[6] Soluble organics are composed mainly of polyphenolics and lower molecular weight polymers (fulvic acids), originating from the decomposition processes of nutrient-poor, acidic organic residues, such as heath and coniferous litter. According to the traditional metal–fulvate theory, organometallic complexes form in the decomposing litter under conditions of low metal saturation, complex more iron (Fe) and aluminum (Al) as they flush down the soil profile with percolating water, and precipitate when the ratio of metal to C reaches a critical level. Other proposed


Fig. 1 Scanning electron micrographs of a fresh (top) and decomposing (bottom) manzanita leaf (magnification: 1000).

1174


mechanisms for immobilization include flocculation, polymerization, and sorption on mineral surfaces.

ENVIRONMENTAL CONTROLS Climate Climate influences organic matter accumulation by controlling the balance between litter production and decomposition rates. On a global basis, soil carbon content increases with increasing precipitation but decreases with increasing temperature.[7] In situations in which moisture is not a controlling factor, decomposition increases with increasing temperature. Litter production follows the same trend. Therefore, the worldwide accumulation of soil organic matter is related more to factors controlling decomposition than to the productivity of ecosystems. On a local and regional scale, OM content decreases exponentially with rising annual temperature at any given level of precipitation:[1] OM ¼ CekT

ð2Þ

where T is the mean annual air temperature ( C), and C and k are constants. For instance, in North American prairie soils, carbon content decreases two to three times for each 10 C increase in mean annual temperature when other factors are kept constant. On the other hand, organic matter accumulation does not follow a climatic pattern in poorly drained soils where anaerobic conditions impede decomposition processes. In addition to its influence on total C content, climate may also affect the chemical composition of soil organic matter. High rainfall favors high leaching regimes and reduces the development of aromaticity in soil organic matter.[4] Aromaticity also has been reported to be negatively correlated to the precipitation=temperature ratio. For prairie soils of the Great Plains, polysaccharides decrease with increasing temperature but increase with increasing precipitation.[8]

horizon. Coniferous litter, such as pine, tends to be low in nutrients but high in recalcitrant constituents, such as lignin and waxes, and typically decomposes at slower rates than deciduous litter. The chemical composition of litter also exerts a significant influence on the accumulation and turnover of soil organic matter by determining the palatability of the plant material, which in turn can alter the distribution and activity of soil fauna. Soil animals, such as earthworms, may accelerate decomposition rates by contributing to the rapid mixing of fresh plant residues into the mineral soil. Parent Material Parent material may influence organic matter accumulation through its effect on soil fertility. Soils formed from base cation-rich volcanic rocks (e.g., basalt) are typically more fertile, and thus experience more organic matter accumulation than soils with lower inherent mineral-derived nutrients, such as those formed from granitic materials. Parent material is also effective through its determination of soil texture. Soil clay content and organic matter accumulation are positively correlated. Clay content affects soil moisture and water availability, thereby modifying plant productivity and litter production. Additionally, a high clay content may induce organic matter accumulation by stabilizing humic substances formed during decomposition. Clay and organic matter form organomineral complexes that are resistant to further biodegradation. Topography Topography interacts with microclimate to influence organic matter distribution in soils. Organic matter accumulation is often favored at the bottom of hills where conditions are wetter than at mid- or upperslope positions. In a similar fashion, organic matter accumulation is usually greater on north-facing slopes compared with south-facing slopes (in the Northern Hemisphere) because temperature is lower.

Organisms Time The amount, placement, and chemical composition of the organic residues of vegetation also affects organic matter accumulation.[9] Litter production worldwide declines with increasing latitude from tropical to arctic forests, following the same distribution patterns as net primary productivity. The placement of organic residues affects the distribution of organic matter accumulation within the soil profile. In grassland soils, where belowground production is abundant, organic matter is more evenly distributed than it is in forest soils, where most accumulation occurs in the uppermost


Long-term rates of organic matter accumulation in Holocene-aged soils vary from about 1 to 12 g C m2 yr1.[10] Organic matter, however, does not accumulate indefinitely in soils. Depending on other soil forming factors, an equilibrium level is reached over time. Organic matter encompasses a series of pools with varying turnover rates. Amounts of the young, labile organic matter may level off in decades because plant biomass stabilizes, while amounts of the more recalcitrant fractions, composed of humic


substances often complexed with clay minerals, may continue to increase for tens of thousands of years.

REFERENCES 1. Jenny, H. Factors of Soil Formation; McGraw-Hill: New York, 1941. 2. Buol, S.W.; Hole, F.D.; McCracken, R.J. Pedogenic processes: Internal, soil-building processes. In Soil Genesis and Classification, 3rd Ed.; Iowa State University Press: Ames, 1989; 114–125. 3. Stevenson, F.J. Biochemistry of the formation of humic substances. In Humus Chemistry: Genesis, Composition, Reactions, 2nd Ed.; John Wiley and Sons: New York, 1994; 188–211. 4. Preston, C.M. Applications of NMR to soil organic matter analysis: History and prospects. Soil Science 1996, 161 (3), 145–166.


1175

5. Oades, J.M. The retention of organic matter in soils. Biogeochemistry 1988, 5, 35–70. 6. Browne, B.A. Towards a new theory of podzolization. In Carbon Forms and Functions in Forest Soils; McFee, W.W., Kelly, J.M., Eds.; Soil Science Society of America: Madison, Wisconsin, 1995; 253–273. 7. Post, W.A.; Emanuel, W.R.; Zinke, P.J.; Stangenberger, A.G. Soil carbon pools and world life zones. Nature 1982, 298 (8), 156–159. 8. Amelung, W.; Flach, K.; Zech, W. Climatic effects on soil organic matter composition in the great plains. Soil Sci. Soc. Am. J. 1997, 61, 115–123. 9. Quideau, S.A.; Anderson, M.A.; Graham, R.C.; Chadwick, O.A.; Trumbore, S.E. Soil organic matter processes: characterization by 13C NMR and 14C Measurements. Forest Ecology and Management 2000, 138, 19–27. 10. Chadwick, O.A.; Kelly, E.F.; Merritts, D.M.; Amundson, R.G. Carbon dioxide consumption during soil development. Biogeochemistry 1994, 24, 115–127.

Organic Matter and Global C Cycle Keith Paustian Colorado State University, Fort Collins, Colorado, U.S.A.

INTRODUCTION

CARBON STOCKS IN SOIL

Soil organic matter (SOM) comprises an integral component of the global carbon (C) cycle, both as the largest overall repository of C within the terrestrial system and as a major source and sink for C exchanges between the atmosphere, terrestrial vegetation, and aquatic environments. Consequently, SOM plays a significant role in regulating the composition of the atmosphere, particularly with respect to carbon dioxide (CO2). Thus, there is both concern that the effects of climate and land use change on soils will exacerbate the problem of increasing CO2 in the atmosphere and hope that through better management, soils can play a part in mitigating increasing CO2 levels.

Carbon comprises a relatively minor component of most soils, in terms of mass. Most soils contain from 1% to 10% C by mass in surface horizons, with the majority having C contents in the range of 1–3%. The most notable exceptions are soils formed under waterlogged conditions, which restricts the flow of oxygen to soil organisms, greatly reducing the rates of SOM decay. Such organic soils or ‘‘histosols’’ may contain 10–30% or more of their total dry mass as C. Despite this generally low concentration, on an area basis, the C contained in soils usually exceeds that contained in the living and dead vegetation (Table 1). Estimates of the global C stock in soils vary, although most recent estimates are on the order of 1400–1600 Pg (Petagram ¼ 1015 g ¼ billion metric tonnes) organic C[2,3] and 700–900 Pg inorganic C.[4] Concentrations of organic C are highest in wetlands and in cold, wet environments (e.g., boreal forest) where decomposition rates are suppressed, and are lowest in desert and tundra soils where plant productivity, and hence C additions to soil, is low. Stocks of inorganic C are greatest in desert and semiarid to subhumid grasslands and savannas, where carbonate leaching is restricted and the formation of secondary carbonate minerals is favored.

CARBON COMPONENTS OF SOIL Carbon is present in both inorganic and organic forms in soil. With the exception of some soils formed on carbonate-rich parent material or arid soils containing high levels of primary or secondary carbonates, organic forms of C predominate in soil. These organic compounds range from fresh plant residues— which are the primary source of SOM—to highly recalcitrant, amorphous humic substances. Plant residues are broken down by the soil biota, chiefly microorganisms, to derive energy (i.e., respiration) and C and nutrient elements needed to grow and reproduce. The soil biota comprise a complex food web in which microbial-, plant-, and animal-derived materials are continually consumed, decomposed, and reformed as soil biomass and other secondary compounds. Plant compounds that are not fully metabolized by the biota, e.g., lignin derivatives, microbial metabolites and other organic substances, can also recombine through chemical and physical reactions. Thus, the decomposition process can be viewed as a ‘‘cascade’’ of biophysiochemical reactions and products,[1] resulting in the loss of C via respiration, together with the formation of a wide array of secondary SOM compounds including complex recalcitrant substances which can persist in soil for several hundreds to thousands of years. 1176 Copyright © 2006 by Taylor & Francis

CARBON FLUXES The CO2 exchanges between the atmosphere, vegetation, and soils are among the largest annual fluxes in the global C cycle (Fig. 1), which are clearly detectable in the regular, seasonal variations in atmospheric CO2 concentrations. Global estimates of net primary production, expressed as C assimilated by live plants (minus respiration) are on the order of 60 Pg C=yr.[5] Much of the annually produced biomass is added to the soil and soil-surface each year as senesced leaves, roots, and woody debris. Emissions of CO2 from soils through decomposition processes (via heterotrophic respiration) are thought to be on the order of 50 Pg=yr with additional losses of about 10 Pg=yr from disturbances such as fire.[5] If terrestrial biome C stocks were at equilibrium, then additions of C through net Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001824 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter and Global C Cycle

1177

Table 1 Global carbon stocks in soil and vegetation by major biome types Area (106 km2)

Average C density (Mg ha1)

Soil C stock (Pg)

Vegetation C stock (Pg)

Tropical forests

17.6

Temperate forests

10.4

123

216

212

96

100

59

Boreal forests Tropical savannas

13.7

344

471

88

22.5

117

264

66

Temperate grasslands

12.5

236

295

9

Deserts and semi-deserts

45.5

42

191

8

Biome

Tundra

9.5

13

121

6

Wetlands

3.5

642

225

15

16.0

80

128

3

—

2011

466

Croplands Total

151.2


primary production would be fully balanced by losses through decomposition and other emission sources such as fire. However, overall estimates of the global C cycle suggest that a net accumulation of C presently occurs in terrestrial vegetation and soils, at a rate of about 2 Pg C=yr (Table 2). During the 1980s, tropical regions are believed to have been a net source of C emissions, largely through deforestation. Similar global estimates for the 1990s have not yet been reported, but if the tropics remain a net source of C from deforestation then the size of the present-day terrestrial sink outside the tropics would be correspondingly greater. In the global budget, the terrestrial sink term is calculated as a difference between the estimates of fossil C emissions, ocean uptake, and observed increases in

Fig. 1 Simplified depiction of the global carbon cycle. Total stocks of carbon in the atmosphere (as CO2), oceans, vegetation and soils þ detritus are shown in bold and approximate annual gross fluxes between the major biosphere pools are shown in italics. Dashed arrows denote direction and magnitude of the net flux of CO2 from industrial sources (mainly fossil fuel combustion) and net sinks to terrestrial and marine environments, estimated as averages for 1990–1999. All stock values are in units of Pg (i.e., billion metric tonnes) and fluxes are Pg=yr. (Adapted from Refs.[5,14,24].)


atmospheric CO2. The existence and relative size of the inferred terrestrial sink are broadly consistent with estimates of a northern hemisphere terrestrial sink, based on atmospheric transport models[6]—although considerable uncertainty as to the magnitude and geographic distribution of the sink remains.[7]

SOIL CARBON FLUXES AND GLOBAL CHANGE While the net terrestrial accumulation of C is small relative to the annual fluxes of C between the atmosphere and land, it is extremely significant in relation to the net change in CO2 in the atmosphere. In other words, if the terrestrial C sink was to disappear, and everything else remained equal, the rate of growth of CO2 in the atmosphere would increase by more than 50%, from about 3.2 to over 5 Pg=yr (Table 2). The origin of and controls on this terrestrial C sink are not well understood. Model estimates[8] attribute much of the sink to increased plant growth rates due to higher CO2 concentrations, although other factors such as N deposition and changing land use contribute as well. The contribution of soils to the overall C sink is still uncertain although it is hypothesized that most of the C accumulates as biomass and surface litter pools. In part, this is to be expected due to the lag time between C accumulation in biomass and fresh residue pools, and its subsequent appearance in more stabilized SOM pools, as well as the fact that much of the C added to soils is relatively rapidly decomposed and evolved as CO2. Global accumulation of refractory humic compounds in mature native ecosystems has been estimated at 0.4 Pg=yr and it has been suggested that these increases roughly balance leaching losses and long-term average organic C transfers to oceans.[9] Soil C increases attributed to global afforestation and

1178


Table 2 Major net sources and sinks of C in the global budget for the 1980s and 1990s 1980–1989

1990–1999

Atmospheric increase in CO2-C

3.3 0.1

3.2 0.1

Emission of C from fossil fuel and cement

5.4 0.3

6.3 0.3

Ocean–atmosphere flux

1.9 0.6

1.7 0.3

Land–atmosphere flux (Net from land use change) (Net from other terrestrial sinks)

0.2 0.7 [1.7 (0.6–2.5)]

1.4 0.3 NAa

[1.9 (3.8–0.3)]

NA

a

Not available. (Adapted from Ref.[24].)

However, this legacy of past land use offers an opportunity to reverse the historical trend and through better management of soils, exploit their potential to become a significant sink for CO2. A variety of management options are available to increase soil C storage in agricultural[23] and other managed ecosystems.[5] Recent estimates suggest that potential C sequestration through improvements in land use and management, globally, is on the order of 1 Pg=yr.[5] Thus the role of SOM in the global C cycle, and how it will respond to future changes in climate and land use, will be determined by both the natural forces regulating the Earth’s biosphere as well as the social, economic, and political actions of human kind.

REFERENCES forest regrowth since the 1950s are estimated to be on the order of 0.1 Pg=yr.[10] Global estimates are still lacking for other ecosystems, but recent inventories for the U.S.[11] suggest that soils in cropland and grazing land currently represent a small sink, on the order of 10–30 Tg=yr (Teragram ¼ 1012 g ¼ million metric tonnes), which can be compared to estimates for C increases in U.S. forest biomass of about 210 Tg=yr.[12] The potential feedbacks of climate change on soil C stocks are currently being debated. Since both plant production (hence, C inputs) and decomposition (hence, C losses) are affected by changes in temperature, precipitation, and CO2 concentrations, the interactions and feedbacks controlling the terrestrial C balance are complex and difficult to predict. Earlier estimates, assuming a stimulation of decomposition due to projected increases in temperature, suggested that soils could become a significant net source of CO2, on the order of 40–60 Pg over a 50–100 yr period, under a climate regime predicted for double presentday CO2 concentrations.[13,14] Other studies suggest that including positive effects of CO2 on plant productivity, management, together with adaptations in land might largely offset climate-driven increases in decomposition potential.[15,16] Recent debates[17–19] have highlighted the complexity and continuing uncertainty of how temperature increase, and other changes in climate and CO2, may impact soils and the global C cycle. The other significant driver of soil C changes, both past and future, is land use and management. It is well established that the conversion of native ecosystems (e.g., forests, grasslands, wetlands), primarily to agricultural uses, has led to significant losses of C from terrestrial ecosystems, on the order of 120–170 Pg C or more over the past 150–300 yr from vegetation and soils combined.[20] From soils alone, estimates of historical losses over the same period are 50–100 Pg C.[21,22]


1. Swift, M.J.; Heal, O.W.; Anderson, J.M. Decomposition in Terrestrial Ecosystems; Blackwell Press: Oxford, UK, 1979; 372 pp. 2. Post, W.M.; Emanuel, W.R.; Zinke, P.J.; Stangenberger, A.G. Soil carbon pools and world life zones. Nature 1982, 298, 156–159. 3. Eswaran, H.; Van Den Berg, E.; Reich, P. Organic carbon in soils of the world. Soil Sci. Soc. Am. J. 1993, 57, 192–194. 4. Eswaran, H.; Reich, P.F.; Kimble, J.M.; Beinroth, F.H.; Padmanabhan, E.; Moncharoen, P. Global carbon stocks. In Global Climate Change and Pedogenic Carbonates; Lal, R., Kimble, J.M., Eswaran, H., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 2000; 15–25. 5. IPCC. Land Use, Land Use Change, and Forestry, Intergovernmental Panel on Climate Change Special Report; Cambridge University Press: Cambridge, UK, 2000; 377 pp. 6. Tans, P.P.; Fung, I.Y.; Takahashi, T. Observational constraints on the global atmospheric carbon dioxide budget. Science 1990, 247, 1431–1438. 7. Field, C.B.; Fung, I.Y. The not-so-big US carbon sink. Science 1999, 285, 544–545. 8. McGuire, A.D.; Sitch, S.; Clein, J.S.; Dargaville, R.; Esser, G.; Foley, J.; Heimann, M.; Joos, F.; Kaplan, J.; Kicklighter, D.W.; Meier, R.A.; Melillo, J.M.; Moore, B., III; Prentice, I.C.; Ramankutty, N.; Reichenau, T.; Schloss, A.; Tian, H.; Williams, L.J.; Wittenberg, U. Carbon balance of the terrestrial biosphere in the twentieth century: analysis of CO2, climate and land use effects with four process-based ecosystem models. Glob. Biogeochem. Cycles 2001, 15, 183–206. 9. Schlesinger, W.H. Evidence from chronosequence studies for a low carbon-storage potential of soils. Nature 1990, 348, 232–234. 10. Post, W.M.; Kwon, K.C. Soil carbon sequestration and land-use change: processes and potential. Global Change Biol. 2000, 6, 317–327. 11. Eve, M.D.; Paustian, K.; Follett, R.; Elliott, E.T. A national inventory of changes in soil carbon from national resources inventory data. In Methods of


12.

13.

14.

15.

16.

17.

18.

Assessment of Soil Carbon; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 2001; 593–610. US Environmental Protection Agency. Inventory of US Greenhouse Gas Emissions and Sinks: 1990–1998; US EPA 236-R-00-001; US EPA: Washington, DC, 2000. Jenkinson, D.S.; Adams, D.E.; Wild, A. Model estimates of CO2 emissions from soil in response to global warming. Nature 1991, 351, 304–306. Schlesinger, W.H. An overview of the carbon cycle. In Advances in Soil Science: Soils and Global Change; Lal, R., Kimble, J., Levine, E., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1995; 9–25. Prentice, K.C.; Fung, I.Y. The sensitivity of terrestrial carbon storage to climate change. Nature 1990, 346, 48–51. Paustian, K.; Elliott, E.T.; Peterson, G.A.; Killian, K. Modelling climate, CO2 and management impacts on soil carbon in semi-arid agroecosystems. Plant Soil 1996, 187, 351–365. Giardina, C.P.; Ryan, M.G. Evidence that decomposition rates of organic carbon in forest mineral soil do not vary with temperature. Nature 2000, 404, 858–861. Davidson, E.A.; Trumbore, S.E.; Amundson, R. Soil warming and organic carbon content. Nature 2000, 408, 789–790.


1179

19. Thornley, J.H.M.; Cannell, M.G.R. Soil carbon storage response to temperature: a hypothesis. Ann. Bot. 2001, 87, 591–598. 20. Houghton, R.A. The annual net flux of carbon to the atmosphere from changes in land use 1850–1990. Tellus 1999, 51B, 298–313. 21. IPCC. Climate change 1995: chapter 23 – agricultural options for mitigation of greenhouse gas emission impacts. In Adaptations and Mitigation of Climate Change: Scientific–Technical Analyses; Intergovernmental Panel on Climate Change (IPCC) Working Group II; Cambridge University Press: Cambridge, UK, 1996; 745–771. 22. Lal, R. Soil Management and restoration for carbon sequestration to mitigate the accelerated greenhouse effect. Prog. Environ. Sci. 1999, 1, 307–326. 23. Paustian, K.; Collins, H.P.; Paul, E.A. Management controls on soil carbon. In Soil Organic Matter in Temperate Agroecosystems; Long-Term Experiments of North America; Paul, E.A., Paustian, K., Elliott, E.T., Cole, C.V., Eds.; CRC/Lewis Publishers: Boca Raton, FL, 1997; 15–49. 24. IPCC. Climate change 2001: the scientific basis. In Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change (IPCC); Houghton, J.T., Yihui, D., Eds.; Cambridge University Press: Cambridge, UK.

Organic Matter and Nutrient Dynamics Charles W. Rice Kansas State University, Manhattan, Kansas, U.S.A.

INTRODUCTION Soil organic matter (SOM) is a fundamental component of soil and the global carbon (C) cycle. Soil organic matter controls many of the chemical, physical, and biological properties of the soil.[1] The estimated amount of organic C stored in world soils is about 1100–1600 Pg, more than twice the C in living vegetation (560 Pg) or in the atmosphere (750 Pg).[2] Hence, even relatively small changes in soil C storage per unit area could have a significant impact on the global C balance. Soil organic matter is derived mostly from plant residues. Plants convert CO2 into tissue through photosynthesis. Upon their death, plant tissues decompose, primarily by soil microorganisms, and most of the C in the plant material is eventually released back into the atmosphere as CO2. Between 10% and 20% of the C in plant residue forms SOM, sometimes referred to as ‘‘humus.’’ Some of this C can persist in soils for hundreds and even thousands of years. Associated with the C in soil organic matter are many essential plant nutrients—primarily N, P, and S. Concentrations of soil organic matter range from 0.2 to over 80% in peat soils although the typical range for temperate soils is 0.4–10%.[3] While it is a minor component of most soils, SOM is essential to the living component of soil. Soil organic matter provides the energy source for most soil microorganisms, and provides the nutrient for plants and the soil biological community. Thus, knowledge of SOM dynamics is crucial for the understanding of global C cycling and plant production. The accumulation of SOM is dependent on the quantity and quality of organic residue inputs, largely as plant material, the rates of microbial decomposition, and the capacity of the soil to store organic matter. The quality of the plant residue affects both the extent and rate of decomposition. Labile C compounds such as simple sugars degrade relatively rapidly and more completely to CO2. On the other end of the spectrum, lignin is more difficult to degrade. Most microorganisms do not have the capacity to completely degrade lignin to CO2. Thus, many of the partial degradation products form the precursors to soil organic matter. Generally, the C : N ratio is a guide of decomposability. A ratio >30 slows decomposition and immobilizes N; a ratio 250 mm provide the greatest protection.[5,6]

COMPOSITION Soil organic matter is not one definable entity. Since SOM is formed from plant material and microbial decomposition products, it is a myriad of organic compounds. There have been several theories on the formation of soil humus but the most widely accepted theory is that organic residues undergo decomposition by microorganisms.[7] The altered compounds and new compounds synthesized by soil microbes polymerize through chemical or enzymatic reactions. Thus, SOM is undergoing constant transformation. Typically, most soil organic models define three pools of SOM (Fig. 1). The active pool, which is comprised of microbial biomass and labile organic compounds makes up less than 5% of the soil organic C. The slow pool usually makes up 20–40% of the total organic C and the recalcitrant pool makes up 60–70% of the soil C. These fractions Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120002258 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter and Nutrient Dynamics

Fig. 1 Schematic of plant decomposition through microbial biomass in the formation of soil organic matter. (Adapted from Ref.[7].)

are often defined kinetically based on laboratory mineralization.[5,8] Microbial biomass is the processor and the slow pool is the one in which much of the plant-associated nutrients reside for mineralization. The recalcitrant pool is material that is difficult to degrade and contains what in the older literature was known as humic and fulvic acids—fractions obtained by chemical-fractionation procedures. The active pool has turnover times on the order of months to years, the slow pool takes decades to turn over, while C in the recalcitrant pool takes from hundred to thousands of years to turn over completely. However, 2–5% of the recalcitrant pool is degraded annually. Since the recalcitrant pool is generally in equilibrium in natural systems, then the rate for formation equals the rate of degradation.

1181

N as inorganic N from the soil can be a significant source of N to satisfy plant needs. This process is called net N mineralization, which is the sum of two simultaneous processes, N mineralization and N immobilization. Nitrogen mineralization, or more correctly called ‘‘ammonification,’’ is the conversion of organic N to ammonium. Nitrogen immobilization is the conversion of ammonium to organic N. Microorganisms control both these processes, thus factors that regulate microbial activity, as described earlier, will impact N availability.[9] In addition to environmental factors, the quality of the organic substrate for microorganisms is important. Low-quality substrate, i.e., high C : N ratio, microorganisms degrading the residue require additional nutrients, primarily N. As a result, soil microorganisms assimilate or immobilize inorganic N. Plants may become deficient in N as the microbes fulfill their N need during decomposition of high C : N residue. Later as the organic material is processed, the N previously assimilated by microorganisms is re-released as inorganic N and can become available to plants. Often the release of organic N to the inorganic forms is in synchrony with plant uptake since favorable temperatures and water availability that promote microbial activity also promote plant growth.[9,10] In most native ecosystems and organic agriculture, N mineralization is the major source of plant N needs. In cropland as much as 11–300 kg N ha1 can be supplied from organic matter.[11] Further biological transformations of N occur in the soil, including nitrification (oxidation of NH4þ to NO3) and denitrification (conversion of NO3 to N2); however, these are not directly linked to SOM and will not be discussed in this section. Please refer to Sylvia et al.[12] for further discussion.

NUTRIENTS Along with carbon, SOM contains important plant nutrients. Soil organic matter can be a source or sink of plant nutrients. Plant productivity is directly associated with SOM content and turnover. Approximately 90–95% of the soil N, 40% of the soil P, and 90% of the soil S is associated with SOM.[3] Generally, the C : N : P : S ratio is 100 : 10 : 1 : 1. In agricultural soils, approximately 2–4% of the organic matter is rendered available for plant uptake on an annual basis. As discussed earlier, the more active pools of SOM are likely to be the major source of plant nutrients.

Nitrogen The plant nutrient that is needed from soil in the greatest quantities is N. Approximately 90–95% of the soil N is in the organic form. The net release of organic


Phosphorus and Sulfur Sulfur follows similar transformations as N. P is more controlled by soil chemistry and chemical transformations. Thus, SOM is not usually a major source of P by plants, except in very high organic soils. However, organic P can represent 80% of the total P in some soils. Significant amounts of P and S are contained within the microbial biomass. As a rule, a C : S or C : P ratio greater than 60 promotes immobilization of S and P into the microorganisms. Because of the importance of SOM to the quality of the soil and plant productivity, an understanding of SOM dynamics is critical to preserve natural ecosystems and ensure the long-term productivity of managed ecosystems. Gains and losses of SOM have added significance because it is a reservoir for global C and the associated interaction with climate change.

1182


Table 1 Land use for C sequestration Management strategies Land use

Soil management

Crop management

Cultivation

Tillage

Varieties

Rangeland

Residue management

Crop rotations

Forestry

Fertility Water management Erosion control

Cover crops


MANAGEMENT OF SOM TO ENHANCE SOIL QUALITY

Fig. 2 Relationship between SOM and soil, water, and air quality. (Adapted from Ref.[14].)

Agriculture in the 1800s and early 1900s relied upon the plowing the soil with low crop yields and crop residues were often removed. This combination of agricultural practices resulted in reducing the replenishment of organic C to the soil. Approximately, 50% of the SOM has been lost from soil over a period of 50–100 yr of cultivation. In recent decades, higher yields, return of crop residues, and development of conservation tillage practices have increase SOM. Table 1 lists several practices affecting the soil’s ability to sequester C.[13] Examples of rates of soil C increases are summarized in Table 2.[14] Nitrogen management that increases crop productivity results in an increase in SOM.[15] Nitrogen fertilizer applied at recommended rates for 10 yr increase soil C approximately 2 MT C ha1. Grassland systems also can contribute to C sequestration when properly managed. Under elevated atmospheric CO2, the soil contained 6% more C to a depth of 15 cm compared with ambient conditions.[16] The increase in soil C was due to increased plant production followed by incorporation into the soil. The amount of C sequestered over the 8 yr experimental period was equivalent to 4 Mg ha1. Proper fire management may also increase soil C.[17] Managing agricultural soils for sequestering C will result in additional benefits. Increasing soil organic C include increased crop productivity and enhanced soil,

Table 2 Estimates of C sequestration potential of agricultural practices of U.S. cropland Agricultural practice Conservation Reserve Program

(MTC/ha/yr) 0.3–0.7

Conservation tillage

0.24–0.40

Fertilizer management

0.05–0.15

Rotation with winter cover crops

0.1–0.3

Summer fallow elimination

0.1–0.3

[14]

(Adapted from Ref.

.)


water, and air quality (Fig. 2). In addition, management practices that increase soil C also tend to reduce soil erosion, reduce energy inputs, and improve soil resources.

REFERENCES 1. Doran, J.W.; Parkin, T.B. Defining and assessing soil quality. In Defining Soil Quality for a Sustainable Environment; Doran, J.W., Coleman, D.C., Bezdicek, D.F., Stewart, B.A., Eds.; [Spec. Publ 35]; Soil Sci. Soc. Am.: Madison, WI, 1994; 3–24. 2. Sundquist, E.T. The global carbon dioxide budget. Science 1993, 259, 934–941. 3. Smith, J.L.; Lynch, J.M.; Bezdicel, D.F.; Papendick, R.I. Soil organic matter dynamics and crop residue management. In Soil Microbial Ecology; Metting, B., Ed.; Marcel Dekker: New York, NY, 1992; 65–94. 4. Linn, D.M.; Doran, J.W. Effect of water-filled pore space carbon dioxide and nitrous oxide production in tilled and nontilled soils. Soil Sci. Soc. Am. J. 1984, 48, 1267–1272. 5. van Veen, J.A.; Paul, E.A. Organic C dynamics in grassland soils. 1. Background information and computer simulations. Can. J. Soil Sci. 1981, 61, 185–201. 6. Jastrow, J.D.; Miller, R.M. Soil aggregate stabilization and carbon sequestration: feed backs through organomineral associations. In Soil Process and the Carbon Cycle; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press: New York, 1996; 207–223. 7. Paul, E.A.; Clark, F.E. Soil Biology and Biochemistry; Academic Press: San Diego, CA, 1996. 8. Rice, C.W.; Garcia, F.O. Biologically active pools of soil C and N in tallgrass prairie. In Defining Soil Quality for a Sustainable Environment; Doran, J.W., Coleman, D.C., Bezdicek, D.F., Stewart, B.A., Eds.; [Spec. Publ 35]; Soil Sci. Soc. Am.: Madison, WI, 1994; 201–208.


9. Rice, C.W.; Havlin, J.L. Integrating mineralizable N indices into fertilizer N recommendations. In Soil Testing: Prospects for Improving Nutrient Recommendations; [Spec. Pub. No. 40]; Havlin, J.L., Jacobsen, J.S., Eds.; Soil Sci. Soc. Am.: Madison, WI, 1994; 1–13. 10. McGill, W.B.; Meyers, R.J.K. Controls on dynamics of soil and fertilizer nitrogen. In Soil Fertility and Organic Matter as Critical Components of Production Systems; Follett, R.F., Stewart, J.W.B., Cole, C.V., Eds.; [Spec. Pub. No. 19]; Soil Sci. Soc. Am.: Madison, WI, 1987; 73–99. 11. Smith, J.L.; Paul, E.A. The significance of soil microbial biomass estimations. Soil Biochem. 1990, 6, 357–396. 12. Sylvia, D.M.; Fuhrman, J.F.; Hartel, P.G.; Zuberer, D.A. Principles and Application of Soil Microbiology; Prentice Hall: Upper Saddle River, NJ, 1998. 13. Lal, R.; Kimble, J.R.; Follett, R.F.; Cole, C.V. The Potential of U.S. Cropland to Sequester Carbon and


1183

14.

15.

16.

17.

Mitigate the Greenhouse Effect; Ann Arbor Press: Chelsa, MI, 1998. Lal, R.; Follett, R.F.; Kimble, J.; Cole, C.V. Managing U.S. cropland to sequester carbon in soil. J. Soil Water Conserv. 1999, 54, 374–381. Espinoza, Y. Dynamics and Mechanisms of Stabilization of C and N in Soil. Ph.D. dissertation, Kansas State Univ. Williams, M.A.; Rice, C.W.; Owensby, C.E. Carbon and Nitrogen dynamics and microbial activity in tallgrass prairie exposed to elevated CO2 for 8 years. Plant and Soil 2000, 227, 127–137. Rice, C.W.; Owensby, C.E. Effects of fire and grazing on soil carbon in rangelands. In The Potential of U.S. Grazing Lands to Sequester Carbon and Mitigate the Greenhouse Effect; Follet, R., Ed.; Lewis Publishers: Boca Raton, 2000; 323–342.

Organic Matter in the Landscape Henry H. Janzen B. H. Ellert Agriculture and Agri-Food Canada, Lethbridge, Alberta, Canada

D. W. Anderson Department of Soil Science, University of Saskatchewan, Saskatoon, Saskatchewan, Canada

INTRODUCTION Soil organic matter is composed largely of plant tissues decomposing back to the simple molecules from which they were first formed: carbon dioxide, ammonium, water, and various salts. Although often present in large amounts, organic matter is transient—new litter is continually being added, and organic matter already present is always decomposing. So at any time, the amount present reflects the net balance between additions and losses. Both inputs and losses at a site depend on conditions there. Litter inputs from plant growth and decomposition are both influenced by many factors: temperature, moisture, aeration, nutrients, plant community, and more. And wind and water move organic matter-laden soil about the landscape. So the amount and composition of organic matter differs from place to place in the landscape. Organic matter varies in several dimensions: It varies vertically within the profile, and horizontally across the landscape. It changes across time, often by human influence. How organic matter is distributed over the landscape may be as important as its amount in affecting the way an ecosystem behaves.[1] In this review, we describe how organic matter is distributed in ‘‘natural’’ landscapes, show how human activities can alter that pattern, and illustrate with a few examples how organic matter distribution (and redistribution) can affect ecosystem function. ORGANIC MATTER IN LANDSCAPES UNAFFECTED BY HUMAN ACTIVITY Vertical Distribution In most soil profiles, organic matter (or organic carbon) concentration is highest near the surface, where most plant litter is added, and then declines with increasing depth.[2,3] In grasslands, globally, the surface 0.2 m of soil accounts for 42% of the organic carbon in the first 1 m; in shrublands, the proportion is 33% and in forests, it is 50%.[4] 1184 Copyright © 2006 by Taylor & Francis

The rate of organic matter turnover also changes with depth. Compared to that deeper in the profile, surface soil usually has higher proportions of ‘‘young’’ organic matter from recent inputs of plant litter. Radiocarbon studies show that the mean turnover time or radiocarbon ‘‘age’’ of organic matter usually increases with depth.[5–7] Lateral Distribution The amount of organic matter in the soil at a given spot is the result of a complex interaction, over time, of parent material, climate, vegetation, and topography.[8] Among landscapes, at regional scales, organic matter content is controlled largely by precipitation, temperature, soil texture, and vegetation type.[9,10] But within a landscape, organic matter is also influenced by other factors. Topography, through effects on microclimate and water movement, can produce soils of widely different organic matter within meters (Fig. 1).[9,11] Usually, the amount of organic matter is lowest near summits, and highest in toeslope positions.[12–14] This pattern occurs for various reasons:[13,15,16] higher moisture and nutrient status downslope increase litter production; organic matter in lower slope positions may decompose more slowly because of soil conditions (e.g., reduced aeration or accumulated clay); variations in microclimate produce different plant communities across topographical gradients; and erosion may move organic matter downhill. Apart from topography, localized variations in texture, soil chemistry, vegetation,[1,17] and other properties may also cause organic matter to vary within landscapes. Even in landscapes that appear uniform, therefore, soil organic matter content varies significantly over scales of several meters. Tiessen and Santos[18] observed coefficients of variation greater than 50% in organic carbon and total nitrogen concentrations in the surface soil of a tropical semiarid field (65 40 m2) immediately after clearing. Organic matter varies across landscapes not only in amount but also in form. For example, Paul et al.[5] observed that radiocarbon ‘‘age’’ of surface soil Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120002259 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter in the Landscape

1185

Converting grass- and forestlands to arable agriculture, for example, typically results in the loss of about 30% of the organic carbon originally present in the solum.[23] But with better land management, at least part of the organic matter lost can be restored.[24] Human influence on soil organic matter—whether it leads to losses or gains—is rarely uniform over the landscape. Thus, human intervention often alters not only the amount of organic matter, but also its distribution over the landscape. The rearrangement can occur in several ways. Patchwork Application of Practices

Fig. 1 An illustration of variability in soil organic carbon in transects across two toposequences, a native grassland and an agricultural field cultivated for about two decades. Organic matter is about 58% carbon, by weight, so it is often measured as organic carbon. (Adapted from Ref.[28].)

increased from toeslope to summit positions at two uncultivated grassland sites. And Schimel et al.[15] found that the relative mineralization rate of organic matter (N mineralized per unit of total nitrogen) decreased downslope, even though total mineralization increased, pointing to differences in organic matter quality. Temporal Distribution Organic matter in the landscape also changes with time. Soils in early stages of development accumulate organic matter, but the rate of build-up slows as soils approach a steady state, where decomposition roughly balances litter inputs.[19] Even then, however, organic matter fluctuates during a year because litter additions follow a different pattern from decomposition. It also fluctuates from year to year with changes in weather: in some periods, when litter inputs exceed losses, organic matter accumulates; in others, it is depleted.[20] HUMAN INFLUENCE ON DISTRIBUTION Human activities can profoundly alter soil organic matter content. Historically, disturbance of ‘‘natural’’ ecosystems has almost always resulted in losses.[21,22]


Land practices are often not applied uniformly over the landscape. Forests may be cut in patches; diverse cropping practices may be used in scattered patterns in agricultural landscapes; organic materials such as manures, derived from large, far-flung areas, may be funneled into small areas;[25] grazing animals may congregate near water or shade, depositing organic matter there.[26] These, and other practices applied nonuniformly, can redistribute organic matter on the landscape. Nonuniform Effects on C Balance Even where the same management practice or land use change is applied uniformly over a large area, its effect on organic matter may vary from place to place because the landscape is not uniform to begin with. For example, Schimel, Coleman, and Horton[27] observed that proportional losses of organic matter after cultivating grassland were higher in upper- than in lower-slope positions. And the mechanism of organic matter loss, whether by erosion or biological mineralization, may also vary among slope positions.[28] Tillage Farmers cultivate soils to control pests, prepare land for seeding, and bury residues. This tillage dilutes organic matter-rich soil near the surface by mixing it with soil from deeper layers.[14] It may also affect organic matter deeper in the profile by altering rooting patterns, leaching, faunal activity, and soil temperature.[29,30] Consequently, tillage changes organic matter distribution in the profile, even though total amount may not always be affected.[31] Tillage also alters lateral distribution of organic matter by physically ‘‘dragging’’ topsoil. Each tillage pass moves soil, but the extent of movement is greater downslope than upslope, so that over the years, tillage moves soil and organic matter to lower-slope positions.[32]

1186

Erosion Globally, the predominant human influence on distribution of organic matter in the landscape is via erosion, especially in deforested and agricultural lands. Although erosion is a natural process and occurs also in undisturbed ecosystems, human disturbance may increase erosion by orders of magnitude.[33,34] Erosion affects organic matter at a site within the landscape in three ways;[11] it may: remove organic matter; result in mixing of subsoil into the surface layer by stripping surface soil away; or result in deposition of soil eroded from elsewhere in the landscape. The effects of erosion are not uniform across the landscape. Highest losses usually occur from convex uplands or ‘‘shoulder elements.’’[11,13] Much of the eroded soil removed from one location in the landscape may be deposited nearby, especially, in closed watersheds.[35] But areas of net removal are usually larger than areas of deposition so that erosion often increases the variability of organic matter on the landscape. This is compounded by selective translocation of soil fractions rich in organic matter.[14,36] The net effect, therefore, is often a removal of organic matter from widespread areas in the landscape and its deposition in low-lying areas. Erosion not only affects directly the distribution of organic matter on the landscape, but may also have long-lasting secondary consequences through effects on plant growth and litter input.[14] While erosion can result in extensive translocation of organic matter, its net effect on total amount stored in the landscape is not always clear. If erosion suppresses productivity, thereby limiting replenishment of organic matter, the organic matter may spiral downwards over the long term. But when productivity of eroded areas can be restored, the eroded landscape might eventually contain more organic matter than before, because of the higher storage potential of eroded areas[37] and the accumulated and buried organic matter in depositional areas.[14] Disruption and management of ecosystems by humans often alter irreversibly the amount and distribution of organic matter in landscapes, often increasing variability on the landscape (e.g., Fig. 1).[38] Our understanding of this redistribution is still incomplete.[39,40]

IMPLICATIONS The heterogeneity of organic matter in the landscape, and its further rearrangement by human activity, determine how ecosystems function. To illustrate, we present three examples. First, the spatial variability of organic matter influences productivity on the landscape. Plant productivity



is closely linked to organic matter;[41] consequently, landscapes with variable organic matter usually show corresponding heterogeneity in productivity (whether high organic matter increases productivity or high productivity increases organic matter is not always clear). In farmlands, therefore, increasing attention has been devoted recently to ‘‘site-specific farming’’—adjusting practices spatially to compensate for variability in organic matter and related soil properties.[42] Secondly, the way organic matter is distributed across the landscape influences the unintended release of nutrients into the broader environment. For example, organic matter affects transformations of nitrogen, both as a source of mineralized nitrogen and by effects on microbial activity. Consequently, nitrate leaching or nitrous oxide emissions may be linked to organic matter distribution, especially since sites where organic matter accumulates may also have high moisture.[43,44] Thirdly, the heterogeneity of organic matter within the landscape makes it harder to measure changes in carbon storage. Soil organic matter has been proposed as a potential ‘‘sink’’ for carbon; widespread adoption of practices that build organic matter could increase carbon storage in soils, mitigating the increases in atmospheric CO2 linked to global warming.[45] But to quantify that ‘‘sink’’ precisely, organic matter distribution across the landscape would have to be taken into account; differences in organic matter among points on the landscape are usually much greater than the expected response to new management.[46] At one location, Garten and Wullschleger[47] estimated that more than 100 samples would need to be taken to detect a change in soil organic carbon of about 1 Mg C ha1. The redistribution of organic matter by erosion makes measuring of the carbon sink even more complicated.[48] When erosion occurs during the measurement interval, it may be hard to distinguish carbon exchanged with the atmosphere from that merely redistributed on the landscape. The topsoil, enriched in organic matter, forms a veneer over the landscape, one that varies from place to place and year to year. The performance and persistence of ecosystems depend on this thin layer. And how that layer varies over the landscape, especially in response to management, therefore has long-lasting effects on how productive and resilient the ecosystem will be. In the past, human activity has often accentuated native variability, removing organic matter from sites where it was already in short supply and depositing it in areas of excess. New approaches may now favor the preservation of organic matter both in amount and distribution across the landscape. We know something about how past management has affected organic matter distribution; now, we need to learn how restorative practices will shape its future patterns across the landscape.


REFERENCES 1. Herrick, J.E.; Wander, M.M. Relationships between soil organic carbon and soil quality in cropped and rangeland soils: the importance of distribution, composition, and soil biological activity. In Soil Processes and the Carbon Cycle; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1997; 405–425. 2. Ajwa, H.A.; Rice, C.W.; Sotomayor, D. Carbon and nitrogen mineralization in tallgrass prairie and agricultural soil profiles. Soil Sci. Soc. Am. J. 1998, 62, 942–951. 3. Zhao, Q.; Zhong, L.; Yingfei, X. Organic carbon storage in soils of southeast China. Nutr. Cycling Agroecosyst. 1997, 49, 229–234. 4. Jobba´gy, E.; Jackson, R.B. The vertical distribution of soil organic carbon and Its relation to climate and vegetation. Ecol. Appl. 2000, 10 (2), 423–436. 5. Paul, E.A.; Follett, R.F.; Leavitt, S.W.; Halvorson, A.; Peterson, G.A.; Lyon, D.J. Radiocarbon dating for determination of soil organic matter pool sizes and dynamics. Soil Sci. Soc. Am. J. 1997, 61, 1058–1067. 6. Trumbore, S. Age of soil organic matter and soil respiration: radiocarbon constraints on belowground C dynamics. Ecol. Appl. 2000, 10 (2), 399–411. 7. Scharpenseel, H.W.; Pfeiffer, E.M.; Becker-Heidmann, P. Ecozone and soil profile screening for C-residence time, rejuvenation, bomb 14C photosynthetic d13C Changes. In Assessment Methods for Soil Carbon; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers=CRC Press: Boca Raton, FL, 2001; 207–220. 8. Jenny, H. The Soil Resource: Origin and Behavior; Springer: New York, 1980; 377 pp. 9. Arrouays, D.; Daroussin, J.; Kicin, J.L.; Hassika, P. Improving topsoil carbon storage prediction using a digital elevation model in temperate forest soils of france. Soil Sci. 1998, 163 (2), 103–108. 10. Burke, I.C.; Yonker, C.M.; Parton, W.J.; Cole, C.V.; Flach, K.; Schimel, D.S. Texture, climate, and cultivation effects on soil organic matter content in U.S. grassland soils. Soil Sci. Soc. Am. J. 1989, 53, 800–805. 11. Pennock, D.J. Effects of soil redistribution on soil quality: Pedon, landscape, and regional scales. In Soil Quality for Crop Production and Ecosystem Health; Gregorich, E.G., Carter, M.R., Eds.; Elsevier: Amsterdam, 1997; 167–185. 12. Burke, I.C.; Elliott, E.T.; Cole, C.V. Influence of macroclimate, landscape position, and management on soil organic matter in agroecosystems. Ecol. Appl. 1995, 5 (1), 124–131. 13. Schimel, D.S.; Kelly, E.F.; Yonker, C.; Aguilar, R.; Heil, R.D. Effects of erosional processes on nutrient cycling in Semiarid landscapes. In Planetary Ecology; Caldwell, D.E., Brierley, J.A., Brierley, C.L., Eds.; Van Nostrand Reinhold: New York, 1985; 571–580. 14. Gregorich, E.G.; Greer, K.J.; Anderson, D.W.; Liang, B.C. Carbon distribution and losses: erosion and deposition effects. Soil Till. Res. 1998, 47, 291–302.


1187

15. Schimel, D.; Stillwell, M.A.; Woodmansee, R.G. Biogeochemistry of C, N, and P in a soil catena of the shortgrass steppe. Ecology 1985, 66 (1), 276–282. 16. Cheng, W.; Virginia, R.A.; Oberbauer, S.F.; Gillespie, C.T.; Reynolds, J.F.; Tenhunen, J.D. Soil nitrogen, microbial biomass, and respiration along an arctic toposequence. Soil Sci. Soc. Am. J. 1998, 62, 654–662. 17. Mueller-Harvey, I.; Juo, A.S.R.; Wild, A. Soil organic C, N, S and P after forest clearance in nigeria: mineralization rates and spatial variability. J. Soil Sci. 1985, 36, 585–591. 18. Tiessen, H.; Santos, M.C.D. Variability of C, N, and P content of a tropical semiarid soil as affected by soil genesis, erosion and land clearing. Plant Soil 1989, 119, 337–341. 19. Chadwick, O.A.; Kelly, E.F.; Merritts, D.M.; Amundson, R.G. Carbon dioxide consumption during soil development. Biogeochemistry 1994, 24, 115–127. 20. Campbell, C.A.; Zentner, R.P.; Selles, F.; Biederbeck, V.O.; McConkey, B.G.; Blomert, B.; Jefferson, P.G. Quantifying short-term effects of crop rotations on soil organic carbon in southwestern saskatchewan. Can. J. Soil Sci. 2000, 80, 193–202. 21. Solomon, D.; Lehmann, J.; Zech, W. Land use effects on soil organic matter properties of chromic luvisols in semiarid Northern Tanzania: carbon, nitrogen, lignin and carbohydrates. Agric. Ecosyst. Environ. 2000, 78, 203–213. 22. Tiessen, H.; Cuevas, E.; Chacon, P. The role of soil organic matter in sustaining soil fertility. Nature 1994, 371, 783–785. 23. Davidson, E.A.; Ackerman, I.L. Changes in coil carbon inventories following cultivation of previously untilled soil. Biogeochemistry 1993, 20, 161–193. 24. Sampson, R.N.; Scholes, R.J. Additional humaninduced activities—article 3.4. In Land Use, Land Use Change, and Forestry. A Special Report of the IPCC; Watson, R.T., Noble, I.R., Bolin, B., Ravindranath, N.H., Verardo, D.J., Dokken, D.J., Eds.; Cambridge University Press: Cambridge, 2000; 180–281. 25. Fernandes, E.C.M.; Motavalli, P.P.; Castilla, C.; Mukurumbira, L. Management control of soil organic matter dynamics in tropical land use systems. Geoderma 1997, 79, 49–67. 26. Franzluebbers, A.J.; Stuedemann, J.A.; Schomberg, H.H. Spatial distribution of soil carbon and nitrogen pools under grazed tall fescue. Soil Sci. Soc. Am. J. 2000, 64, 635–639. 27. Schimel, D.S.; Coleman, D.C.; Horton, K.A. Soil Organic matter dynamics in paired rangeland and cropland toposequences in North Dakota. Geoderma 1985, 36, 201–214. 28. Gregorich, E.G.; Anderson, D.W. Effects of cultivation and erosion on soils of four toposequences in the Canadian prairies. Geoderma 1985, 36, 343–354. 29. Cihacek, L.J.; Ulmer, M.G. Effects of tillage on profile soil carbon distribution in the northern great plains of the U.S. In Management of Carbon Sequestration in Soil; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1997; 83–91.

1188

30. Mikhailova, E.A.; Bryant, R.B.; Vassenev, I.I.; Schwager, S.J.; Post, C.J. Cultivation effects on soil carbon and nitrogen contents at depth in the Russian chernozem. Soil Sci. Soc. Am. J. 2000, 64, 738–745. 31. Yang, X.-M.; Wander, M.M. Tillage effects on soil organic carbon distribution and storage in a silt loam soil in Illinois. Soil Till. Res. 1999, 52, 1–9. 32. Govers, G.; Lobb, D.A.; Quine, T.A. Tillage erosion and translocation: emergence of a new paradigm in soil erosion research. Soil Till. Res. 1999, 51, 167–174. 33. Meade, R.H.; Yuzyk, T.R.; Day, T.J. Movement and storage of sediment in rivers of the United States and Canada. In The Geology of North America; Wolman, M.G., Riggs, H.C., Eds.; Geological Society of America: Boulder, CO, 1990; Vol. O-1, 255–280. 34. Oldeman, L.R. The global extent of soil degradation. In Soil Resilience and Sustainable Land Use; Greenland, D.J., Szabolcs, I., Eds.; CAB International: Oxfordshire, 1994; 99–118. 35. Pennock, D.J.; De Jong, E. Rates of soil redistribution associated with soil zones and slope classes in Southern Saskatchewan. Can. J. Soil Sci. 1990, 70, 325–334. 36. van Noordwijk, M.; Cerri, C.; Woomer, P.L.; Nugroho, K.; Bernoux, M. Soil carbon dynamics in the humid tropical forest zone. Geoderma 1997, 79, 187–225. 37. Izaurralde, R.C.; Nyborg, M.; Solberg, E.D.; Janzen, H.H.; Arshad, M.A.; Malhi, S.S.; Molina-Ayala, M. Carbon storage in eroded soils after five years of reclamation techniques. In Soil Processes and the Carbon Cycle; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1997; 369–385. 38. Beckett, P.H.T.; Webster, R. Soil variability: a review. Soils Fertil. 1971, 34 (1), 1–15. 39. Starr, G.C.; Lal, R.; Kimble, J.M.; Owens, L. Assessing the impact of erosion on soil organic carbon pools and fluxes. In Assessment Methods for Soil Carbon; Lal, R.,



40.

41.

42.

43.

44.

45.

46.

47.

48.

Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers, CRC Press: Boca Raton, FL, 2001; 417–426. Jacinthe, P.A.; Lal, R.; Kimble, J.M. Assessing water erosion impacts on soil carbon pools and fluxes. In Assessment Methods for Soil Carbon; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers, CRC Press: Boca Raton, FL, 2001; 427–449. Bauer, A.; Black, A.L. Quantification of the effect of soil organic matter content on soil productivity. Soil Sci. Soc. Am. J. 1994, 58, 185–193. Beckie, H.J.; Moulin, A.P.; Pennock, D.J. Strategies for variable rate nitrogen fertilization in hummocky terrain. Can. J. Soil Sci. 1997, 77, 589–595. van Kessel, C.; Pennock, D.J.; Farrell, R.E. Seasonal variations in denitrification and nitrous oxide evolution at the landscape scale. Soil Sci. Soc. Am. J. 1993, 57, 988–995. Pennock, D.J.; Corre, M.D. Development and application of landform segmentation procedures. Soil Till. Res. 2001, 58, 151–162. Lal, R.; Bruce, J.P. The potential of world cropland soils to sequester C and mitigate the greenhouse effect. Environ. Sci. Pol. 1999, 2, 177–185. Ellert, B.H.; Janzen, H.H.; McConkey, B.G. Measuring and comparing soil carbon storage. In Assessment Methods for Soil Carbon; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers/CRC Press: Boca Raton, FL, 2001; 131–146. Garten, C.T., Jr.; Wullschleger, S.D., Jr. Soil carbon inventories under a bioenergy crop (Switchgrass): measurement limitations. J. Environ. Qual. 1999, 2, 1359–1365. Pennock, D.J.; Frick, A.H. The role of field studies in landscape-scale applications of process models: an example of soil redistribution and soil organic carbon modeling using CENTURY. Soil Till. Res. 2001, 58, 183–191.

Organic Matter Management R. Ce´sar Izaurralde Joint Global Charge Research Institute, College Park, Maryland, U.S.A.

Carlos C. Cerri Universidade de Saõ Paulo, CENA, Piracicaba, Saõ Paulo, Brazil

INTRODUCTION Soil organic matter (SOM) consists of a complex array of living organisms such as bacteria and fungi, plant and animal debris in different stages of decomposition, and humus—a rather stable brown to black material showing no resemblance to the organisms from which it originates. Because SOM is or has been part of living tissues, its composition is dominated by carbon (C), hydrogen, oxygen and—in lesser abundance—by nitrogen, phosphorus, sulfur among other elements. Levels of SOM are expressed in terms of soil organic carbon (SOC) concentration (g kg1) or mass per unit area (g m2) to a given depth. The level of SOC in virgin soils reflects the action and interaction of the major factors of soil formation: climate, vegetation, topography, parent material, and age. These factors control SOC content by regulating the balance between C gains via photosynthesis and losses via autotrophic and heterotrophic respiration, as well as C losses in soluble and solid form. The SOC content usually ranges between 5 and 100 g kg1 in mineral soils. These concentrations appear modest but at 1500 Pg, the amount of organic C stored globally in soils is second only to that contained in oceans and at least twice that found in either terrestrial vegetation or the atmosphere. Cultivated soils usually contain less SOC than virgin soils[1] due to the magnification of two biophysical processes: 1) net nutrient mineralization accompanied by release of CO2 due to microbial respiration and 2) soil erosion. SOC losses of up to 50% have been reported within 30–70 years of land use conversions under temperate conditions.[2–5] SOC losses reported in subtropical and tropical environments often match or even surpass those observed under temperate conditions.[6–8] In subtropical and tropical environments, shifting cultivation systems appear to conserve more SOC than forestlands permanently cleared for cultivation.[9]

This manuscript has been created by the Battelle Memorial Institute as operator of the Pacific Northwest National Laboratory under Contract No. DE-AC06–76RLO 1830 with the U.S. Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120002260 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

MAJOR PROCESSES LEADING TO CARBON LOSSES FROM SOIL Mineralization Processes Depending on its frequency and kind, tillage changes the soil biophysical environment in ways that affect the net mineralization of nutrients and the release of carbon. These changes can be described in terms of increases or decreases in soil porosity, disruption of soil aggregates, and redistribution in the proportion of soil aggregate size, as well as alteration of energy and water fluxes. All these changes enhance, at least temporarily, the conversion of organic C into CO2[10] and the net release of nutrients from SOM. Much of the success of past agricultural practices relied heavily on the control of decomposition processes through tillage operations to satisfy plant nutrient demands. All this came at a price, however, for a heavy reliance on soil nutrients to feed crops without proper replenishment led to the worldwide declines of SOM.[11] Soil Erosion Processes Agricultural ecosystems normally experience soil losses at rates considerably greater than natural ecosystems because of an incomplete plant or residue cover of the soil during rainy or windy conditions. When surface and environmental conditions are right (i.e., bare soil, sloping land, intense rain, windy weather), the kinetic energy embedded in wind and water is transferred to soil aggregates causing them to be detached and transported away from their original position across fields or downhill. Besides the physical loss of soil particles and on-site impact on soil productivity, the detachment and transport processes also cause aggregate breakdown, thereby exposing labile C to microbial activity. This aggregate breakdown also facilitates the preferential removal of soil materials comprised mainly of humus and clay or silt fractions. Consequently, waterand windborne sediments become enriched in C with respect to the contributing soil. Carbon enrichment ratios ranging from 3 to 360 have been reported.[12,13] The fate of these C-enriched sediments is not well known, 1189

1190

Organic Matter Management

for while transport and burial of C in eroded sediments may lead to ‘‘sequestration,’’[14] it may also result in part of it being emitted back to the atmosphere as CO2.[15]

Whether or not soil C sequestration practices are widely adopted will depend on their value relative to other C capture and sequestration technologies.

RESTORING SOIL ORGANIC MATTER: THE EMERGING SCIENCE OF SOIL CARBON SEQUESTRATION

Mechanisms of Soil C Sequestration

The Role of Long-Term Field Experiments SOM is an essential attribute of soil quality[16] and has an essential role in soil conservation and sustainable agriculture. Many practices—some involving land use changes—have been shown to increase SOM and thus received considerable attention for their possible role in climate change mitigation.[17–19] Carbon sequestration in managed soils occurs when there is a net removal of atmospheric CO2 because C inputs (nonharvestable net primary productivity) are greater than C outputs (soil respiration, C costs related to fossil fuels and fertilizers). Soil C sequestration has the additional appeal that all its practices conform to principles of sustainable agriculture (e.g., reduced tillage, erosion control, diversified cropping systems, improved soil fertility). Longterm field experiments have been instrumental to increase our understanding of SOM dynamics.[20,21] The first and longest standing experiment was started at Rothamsted, England, by J. B. Lawes and J. H. Gilbert who in 1843 began documenting the impact of nutrient manipulation on crop yields and soil properties.[22] Other experiments were initiated thereafter in America, Europe, and Oceania with the goal of discovering interactions among climate, soil, and management practices. The knowledge that emerged from these experiments has been instrumental for the development and testing of agroecosystem and SOM models.[23]

Recent reviews of experimental results have contributed to organize our understanding of the environmental and management controls of soil C sequestration in grassland[27,28] and agricultural[29,30] ecosystems. The use of C balance, soil fractionation, and isotope techniques have been instrumental to reveal how new C (from crop residues, roots, and organic amendments) enters soil, resides shortly (for a few years) in labile soil fractions, and finally becomes a long-time constituent (for hundreds of years) of recalcitrant organo-mineral complexes.[31] Fig. 1 contrasts young (labile) organic matter fractions extracted from two cultivated soils with and without N fertilizer.[32] The amounts of labile organic matter— fine roots and other organic debris—present in each soil

The Global Importance of Soil C Sequestration There appears to be a significant opportunity for managed ecosystems to act as C sinks. For example, results from inverse modeling experiments suggest that during 1988–1992, terrestrial ecosystems may have been sequestering atmospheric C at rates of 1–2.2 Pg y1.[24] Some of the likely causes include the growth of new forest in previously cultivated land[25] and the ‘‘CO2fertilization effect.’’[26] Globally, agricultural soils have been estimated to have the capacity to sequester C at rates of 0.6 Pg y1[11] during several decades. The realization of this potential C sequestration would not be trivial since it would offset roughly about one-tenth of the current emissions from fossil fuels. In the U.S., annual gains in soil C from improved agricultural practices have been estimated at 0.14 Pg yr1.[25]


Fig. 1 Young organic matter (fine roots and other organic debris) extracted from two Cryoboralfs under cereal cropping for 13 years receiving N at annual rates of 0 (A) and 50 kg ha1 (B). The black material is charcoal. (From Ref.[32].)

Soil organic carbon (kg m2) Region/country

Climate

Soil

Duration

Crop/land use

Treatment

Initial

Moldboard plow No tillage

Final

Depth (cm)

Reference

4.95 5.46

20

[37]

Argentina

Temperate humid

Argiudoll

17

Corn–wheat–soybean

Chaco, Argentina

Subtropical semiarid

Alfisol

20 10 60

Highly restored Moderately restored Highly degraded

7.05 3.10 1.50

20

[6]

Rondonia, Brazil

Tropical humid

Forest Pasture Pasture Pasture Pasture Pasture

4.33 5.85 5.26 5.28 6.56 6.12

50

[38]

5 9 20 41 81

1.74 2.01 2.00 1.59

6

[39]

30

[40]

4.30 5.00 5.30

15

[35]

2.06 2.22

15

[41]

3.52 3.34 3.22

15

[42]

Georgia, U.S.

Temperate humid

Hapludult

5

Bermudagrass

Unharvested Lightly grazed Heavily grazed Hayed

Kentucky

Temperate humid

Paleudalf

20

Conventional tillage corn

0 kg N=ha1 84 kg N=ha1 168 kg N=ha1 336 kg N=ha1

4.89 5.63 5.64 6.14

No till corn

0 kg N=ha1 84 kg N=ha1 168 kg N=ha1 336 kg N=ha1

5.54 5.84 5.89 6.63

Corn

Conventional Organic Manure

No crops—natural succession

Tillage Control

Cont. wheat Fallow–wheat–wheat Green manure–wheat–wheat

Minimum tillage

Kuztown, Pennsylvania

Temperate humid

Fragiudalf

15

Michigan

Cool temperate humid

Hapludalf

7

Swift Current, Canada

Cold semiarid

Haploboroll

10

1.39

4.20 4.40 4.10

3.05 2.99 2.89


1191

(Continued)


Table 1 Examples of worldwide land use and management impacts on SOC

1192

Table 1 Examples of worldwide land use and management impacts on SOC (Continued) Soil organic carbon (kg m2) Region/country Breton, Canada

Climate Cold subhumid

Soil

Duration

Cryoboralf

51

Russia

Cool temperate humid

Mollisol

300 50 100 50

Punjab, India

Subtropical subhumid

Alluvial

6

Morocco

Warm temperate semiarid

Western Nigeria

Treatment

Wheat–fallow

Nil Fertilizer Manure

Wheat–oat–barley–hay–hay

Nil Fertilizer Manure

Initial

Final

Depth (cm)

Reference

2.64

1.81 2.13 3.11

15

[34]

2.07 2.13 1.59 1.51

50

[5]

0.48

15

[43]

3.73 3.39

20

[44]

2.77 2.96 3.90 1.29

15

[45]

2.91 3.37 4.32

Native grassland Hay Continuous cropping Continuous fallow Corn–wheat

11

Continuous wheat and other rotations

20 25 10 10

Bush fallow Bush fallow Bermudagrass Cultivation

Minimum tillage, residue retained Minimum tillage, residue removed Conventional tillage Conventional tillage No tillage

0.48 0.50 3.20



Calcixeroll

Crop/land use


reflect differences in crop productivity induced by addition of N at annual rates of 50 kg ha1 for 13 years. ‘Terra Preta’ soil—in tropical regions of South America and West Africa—represents a prime example of ancient wisdom applied to develop sustainable agriculture through the improvement of soil fertility and SOM.[33] The quantity and quality of C entering soil as well as the interaction of this C with the soil biophysical environment are major factors determining the rate and duration of soil C sequestration. The quantity of C added to soil in the form of roots, crop residues, and organic amendments has been shown to play a dominant role in defining the trajectory of SOC over time.[34] Management practices geared toward optimizing nutrient supply and building nutrient reserves (e.g., fertilization, use of legumes in crop rotations) are almost guaranteed to increase soil C stocks. The quality of crop residues and the timing of their incorporation to soil also have an influential role on C decomposition and, thus, on soil C storage.[35] The degree of soil disturbance—through its impact on soil aggregation—constitutes another major factor regulating C decomposition and retention in soil.[36] In this context, no tillage agriculture has come to represent one of the most significant technological innovations of the last 30 years because it allows farmers the possibility of growing crops economically while reducing erosion and improving both quantity and quality of SOM. A few examples of the management impacts on soil C sequestration from around the world are presented in Table 1. Soil Organic Matter, Energy, and Full C Accounting Land is the natural habitat of humans. Humans dwell on it and use it as a resource for the production of food, fiber, and other goods. Simply put, land is managed when there is a manipulation of energy and matter flows in order to meet certain economic and social objectives. Farm mechanization and fertilizers are two of the many technical innovations that—though they rely on the utilization of fossil energy—have brought dramatic increases in food production during the last century. Changes in management practices that include soil C sequestration as an objective require careful evaluation of their impact not only on soil C gains but also on C costs from the use of fossil energy (e.g., manufacture of fertilizers)[46,47] and on the net greenhouse gas emissions.[48]

THE ROLE OF SOIL ORGANIC MATTER IN THE 21ST CENTURY SOM has played and will continue to play a central role in sustainable land management. The restoration of


1193

SOM at global scales offers a unique opportunity to mitigate global warming. As population levels and affluence increase, demands on land to produce food, fiber, biomass, and other products will remain high. Because land is finite, important decisions will have to be made in order to balance such demands with functional objectives such as the preservation of natural ecosystems. As part of any climate policy, the impact of land use changes and management on SOM storage should be included as a criterion for making these decisions. Depending on their degree of expansion, several evolving agricultural technologies—such as genetically modified crops, conservation tillage, organic farming, and precision farming—may have important implications for soil C sequestration.[19] Their ultimate impact on C sequestration will depend not only on the economic benefit realized by individual producers but also on whether society recognizes the value of soil C storage to mitigate global warming.

REFERENCES 1. Davidson, E.A.; Ackerman, I.L. Changes in soil carbon following cultivation of previously untilled soils. Biogeochemistry 1993, 20, 161–164. 2. Dalal, R.C.; Mayer, R.J. Long-term trends in fertility of soils under continuous cultivation and cereal cropping in southern Queensland. II. Total organic carbon and its rate of loss from the soil profile. Aust. J. Soil Res. 1986, 24, 281–292. 3. Mann, L.K. Changes in soil carbon after cultivation. Soil Sci. 1986, 142, 279–288. 4. Ellert, B.H.; Gregorich, E.G. Storage of carbon, nitrogen and phosphorus in cultivated and adjacent forested soils of Ontario. Soil Sci. 1996, 161, 587–603. 5. Mikhailova, E.A.; Bryant, R.B.; Vassenev, I.I.; Schwager, S.J.; Post, C.J. Cultivation effects on soil carbon and nitrogen contents at depth in the Russian chernozem. Soil Sci. Soc. Am. J. 2000, 64, 738–745. 6. Abril, A.; Bucher, E.H. Overgrazing and soil carbon dynamics in the western chaco of Argentina. Appl. Soil Ecol. 2001, 16, 243–249. 7. Lal, R. Deforestation and land use effects on soil degradation and rehabilitation in western Nigeria, II. Soil chemical properties. Land Degrad. Dev. 1996, 7, 87–98. 8. Lobe, I.; Amelung, W.; Du Preez, C.C. Losses of carbon and nitrogen with prolonged arable cropping from sandy soils of the South African highveld. Eur. J. Soil Sci. 2001, 52, 93–101. 9. Houghton, R.A. Changes in the storage of terrestrial carbon since 1850. In Soils and Global Change; Lal, R., Kimble, J., Levine, E., Stewart, B.A., Eds.; CRC= Lewis Publishers: Boca Raton, FL, 1995; 45–65. 10. Reicoski, D.C.; Lindstrom, M.J. Fall tillage method: effect from short-term carbon dioxide flux from soil. Agron. J. 1993, 85, 1237–1243.

1194

11. Cole, V.; Cerri, C.; Minami, K.; Mosier, A.; Rosenberg, N.J.; Sauerbeck, D. Agricultural options for mitigation of greenhouse gas emissions. In Climate Change 1995: Impacts, Adaptations and Mitigation of Climate Change; Watson, R.T., Zinowera, M.C., Moss, R.H., Eds.; Report of IPCC Working Group II; Cambridge University Press: London, UK, 1996; 745–771. 12. Sterk, G.; Herrmann, L.; Bationo, A. Wind-blown nutrient transport and soil productivity changes in southwest Niger. Land Degrad. Dev. 1996, 7, 325–336. 13. Zobeck, T.M.; Fryrear, D.W. Chemical and physical characteristics of wind-blown sediment. Trans. Am. Soc. Agric. Eng. 1986, 29, 1037–1041. 14. Stallard, R.F. Terrestrial sedimentation and the carbon cycle: coupling weathering and erosion to carbon burial. Global Biogeochem. Cycles 1998, 12, 231–257. 15. Lal, R. Global soil erosion by water and C dynamics. In Soils and Global Change; Lal, R., Kimble, J., Levine, E., Stewart, B.A., Eds.; CRC=Lewis Publishers: Boca Raton, FL, 1995; 131–141. 16. Doran, J.W., Coleman, D.C., Bezdicek, D.F., Stewart, B.A., Eds.; Defining Soil Quality for a Sustainable Environment; Soil Science Society America Special Publication No. 35; SSSA: Madison, WI, 1994; 244 pp. 17. Batjes, N.H. Mitigation of atmospheric CO2 concentrations by increased carbon sequestration in the soil. Biol. Fert. Soils 1998, 27, 230–235. 18. Post, W.M.; Kwon, K.C. Soil carbon sequestration and land-use change: processes and potential. Global Change Biol. 2000, 6, 317–327. 19. Izaurralde, R.C.; Rosenberg, N.J.; Lal, R. Mitigation of climatic change by soil carbon sequestration: issues of science, monitoring and degraded lands. Adv. Agron. 2001, 70, 1–75. 20. Powlson, D.S., Smith, P., Smith, J.U., Eds.; Evaluation of Soil Organic Matter Models Using Existing Long-Term Datasets; NATO ASI Series I; Springer: Heidelberg, 1996; Vol. 38, 429 pp. 21. Paul, E.A., Paustian, K., Elliott, E.T., Cole, C.V., Eds.; Soil Organic Matter in Temperate Agroecosystems: Long-Term Experiments in North America; NATO ASI Series I; CRC=Lewis Publishers: Boca Raton, FL, 1997; 414 pp. 22. Jenkinson, D.S. The rothamsted long-term experiments: are they still of use? Agron. J. 1991, 83, 2–10. 23. Smith, P.; Smith, J.U.; Powlson, D.S.; McGill, W.B.; Arah, J.R.M.; Chertov, O.; Coleman, K.W.; Franko, U.; Frolking, S.; Jenkinson, D.S.; Jensen, L.S.; Kelly, R.H.; Klein-Gunnewiek, H.; Komarov, A.S.; Li, C.; Molina, J.A.E.; Mueller, T.; Parton, W.J.; Thornley, J.H.M.; Whitmore, A.P. A comparison of the performance of nine soil organic matter models using datasets from seven long-term experiments. Geoderma 1997, 81, 153–225. 24. Fan, S.; Gloor, M.; Mahlman, J.; Pacala, S.; Sarmiento, J.; Takahashi, T.; Tans, P. A large terrestrial carbon sink in North America implied by atmospheric and oceanic carbon dioxide data and models. Science 1998, 282, 442–446. 25. Houghton, R.A.; Hackler, J.L.; Lawrence, K.T. The U.S. carbon budget: contributions from land use change. Science 1999, 285, 574–578.



26. Kimball, B.A. Carbon dioxide and agricultural yield: an assemblage and analysis of 430 prior observations. Agron. J. 1983, 75, 779–782. 27. Conant, R.T.; Paustian, K.; Elliott, E.T. Grassland management and conversion into grassland: effects on soil carbon. Ecol. Appl. 2001, 11, 343–355. 28. Scott, N.A.; Tate, K.R.; Ford-Robertson, J.; Giltrap, D.J.; Smith, C.T. Soil carbon storage in plantations and pastures: land-use implications. Tellus 1999, 51B, 326–335. 29. Janzen, H.H.; Campbell, C.A.; Izaurralde, R.C.; Ellert, B.H.; Juma, N.; McGill, W.B.; Zentner, R.P. Management effects on soil c storage on the Canadian prairies. Soil Till. Res. 1998, 47, 181–195. 30. Paustian, K.; Collins, H.P.; Paul, E.A. Management controls on soil carbon. In Soil Organic Matter in Temperate Ecosystems: Long-Term Experiments in North America; Paul, E.A., Paustian, K., Elliott, E.T., Cole, C.V., Eds.; CRC=Lewis Publishers: Boca Raton, FL, 1997; 15–49. 31. Jastrow, J.D. Soil aggregate formation and the accrual of particulate and mineral-associated organic matter. Soil Biol. Biochem. 1996, 28, 665–676. 32. Solberg, E.D.; Nyborg, M.; Izaurralde, R.C.; Malhi, S.S.; Janzen, H.H.; Molina-Ayala, M. Carbon storage in soils under continuous cereal grain cropping: N fertilizer and straw. In Management of Carbon Sequestration in Soil; Lal, R., Kimble, J., Follett, R., Stewart, B.A., Eds.; CRC/Lewis Publishers: Boca Raton, FL, 1998; 235–254. 33. Glaser, B.; Haumaier, L.; Guggenberger, G.; Zech, W. The ‘‘Terra Preta’’ phenomenon: a model for sustainable agriculture in the humid tropics. Naturwissenschaften 2001, 88, 37–41. 34. Izaurralde, R.C.; McGill, W.B.; Robertson, J.A.; Juma, N.G.; Thurston, J.T. Carbon balance of the breton classical plots after half a century. Soil Sci. Soc. Am. J. 2001, 65, 431–441. 35. Drinkwater, L.E.; Wagoner, P.; Sarrantonio, M. Legume-based cropping systems have reduced carbon and nitrogen losses. Nature 1998, 396, 262–265. 36. Six, J.; Elliott, E.T.; Paustian, K. Soil macroaggregate turnover and microaggregate formation: a mechanism for C sequestration under no-tillage agriculture. Soil Biol. Biochem. 2000, 32, 2099–2103. 37. Alvarez, R.; Russo, M.E.; Prystupa, P.; Scheiner, J.D.; Blotta, L. Soil carbon pools under conventional and no-tillage systems in the argentine rolling pampa. Agron. J. 1998, 90, 138–143. 38. Neill, C.; Cerri, C.C.; Melillo, J.M.; Feigl, B.J.; Steudler, P.A.; Moraes, J.F.L.; Piccolo, M.C. Stocks and dynamics of soil carbon following deforestation for pasˆ nia. In Soil Processes and the Carbon ture in Rondo Cycle; Lal, R., Kimble, J., Follett, R., Stewart, B.A., Eds.; CRC=Lewis Publishers: Boca Raton, FL, 1998; 9–28. 39. Franzluebbers, A.J.; Stuedemann, J.A.; Wilkinson, S.R. Bermudagrass management in the southern piedmont USA: I. Soil and surface residue carbon and sulfur. Soil Sci. Soc. Am. J. 2001, 65, 834–841.


40. Ismail, I.; Blevins, R.L.; Frye, W.W. Long-term notillage effects on soil properties and continuous corn yields. Soil Sci. Soc. Am. J. 1994, 58, 193–198. 41. Richter, D.D.; Babbar, L.I.; Huston, M.A.; Jaeger, M. Effects of annual tillage on organic carbon in a finetextured udalf: the importance of root dynamics to soil carbon storage. Soil Sci. 1999, 149, 78–83. 42. Curtin, D.; Wang, H.; Selles, F.; McConkey, B.G.; Campbell, C.A. Tillage effects on carbon fluxes in continuous wheat and fallow–wheat rotations. Soil Sci. Soc. Am. J. 2000, 64, 2080–2086. 43. Ghuman, B.S.; Sur, H.S. Tillage and residue management effects on soil properties and yields of rainfed maize and wheat in a subhumid subtropical climate. Soil Till. Res. 2001, 58, 1–10. 44. Mrabet, R.; Saber, N.; El-Brahli, A.; Lahlou, S.; Bessam, F. Total, particular organic matter and


1195

45.

46. 47.

48.

structural stability of a calcixeroll soil under different wheat rotations and tillage systems in a semiarid area of Morocco. Soil Till. Res. 2001, 57, 225–235. Lal, R. Land use and soil management effects on soil organic matter dynamics on alfisols in western Nigeria. In Soil Processes and the Carbon Cycle; Lal, R., Kimble, J., Follett, R., Stewart, B.A., Eds.; CRC/Lewis Publishers: Boca Raton, FL, 1998; 109–126. Schlesinger, W.H. Carbon and agriculture: carbon sequestration in soils. Science 1999, 284, 2095. Izaurralde, R.C.; McGill, W.B.; Rosenberg, N.J. Carbon cost of applying nitrogen fertilizer. Science 2000, 288, 811–812. Robertson, G.P.; Paul, E.A.; Harwood, R.R. Greenhouse gases in intensive agriculture: contributions of individual gases to the radiative forcing of the atmosphere. Science 2000, 289, 1922–1925.

Organic Matter Modeling Peter Smith School of Biological Sciences, University of Aberdeen, Aberdeen, U.K.

INTRODUCTION There are a number of approaches to modeling soil organic matter (SOM) turnover including 1) processbased multicompartment models; 2) models that consider each fresh addition of plant debris as a separate cohort which decays in a continuous way; and 3) models that account for C and N transfers through various trophic levels in a soil food web. These approaches are described in more detail below.

PROCESS-BASED, MULTICOMPARTMENT SOM MODELS Most models are process-based, i.e., they focus on the processes mediating the movement and transformations of matter or energy and usually assume first order rate kinetics.[1] Early models simulated the SOM as one homogeneous compartment.[2] Some years later two-compartment models were proposed[3,4] and, as computers became more accessible, multicompartment models were developed.[5,6] Of the 33 SOM models currently represented within the Global Change and Terrestrial Ecosystems (GCTE) Soil Organic Matter Network (SOMNET) database,[7–9] 30 are multicompartment, process-based models. Each compartment or SOM pool within a model is characterized by its position in the model’s structure and its decay rate. Decay rates are usually expressed by first-order kinetics with respect to the concentration (C) of the pool dC=dt ¼ kC where t is the time. The rate constant k of first-order kinetics is related to the time required to reduce by half the concentration of the pool ‘‘when there is no input.’’ The pool’s half-life [h ¼ (ln 2)=k], or its turnover time (t ¼ 1=k) are sometimes used instead of k to characterize a pool’s dynamics: the lower the decay rate constant, the higher the half-life, the turnover time, and the stability of the organic pool. The flows of C within most models represent a sequence of carbon going from plant and animal debris to the microbial biomass, and then to soil organic pools of increasing stability. Some models also use 1196 Copyright © 2006 by Taylor & Francis

feedback loops to account for catabolic and anabolic processes and microbial successions. The output flow from an organic pool is usually split. It is directed to a microbial biomass pool, another organic pool, and, under aerobic conditions, to CO2. This split simulates the simultaneous anabolic and catabolic activities and growth of a microbial population feeding on one substrate. Two parameters are required to quantify the split flow. They are often defined by a microbial (utilization) efficiency and stabilization (humification) factor which control the flow of decayed C to the biomass and humus pools, respectively. The sum of the efficiency and humification factors must be inferior to one to account for the release of CO2. A thorough review of the structure and underlying assumptions of different process-based SOM models is available.[6]

COHORT MODELS DESCRIBING DECOMPOSITION AS A CONTINUUM Another approach in modeling SOM turnover is to treat each fresh addition of plant debris into the soil as a cohort.[5] Such models consider one SOM pool that decays with a feedback loop into itself. For example, Q-SOIL[10] is represented by a single rate equation. The SOM pool is divided into an infinite number of components, each characterized by its ‘‘quality’’ with respect to degradability as well as impact on the physiology of the decomposers. The rate equation for the model Q-SOIL represents the dynamics of each SOM component of quality q and is quality dependent. Exact solutions to the rate equations are obtained analytically.[11]

FOOD-WEB MODELS Another type of model simulates C and N transfers through a food web of soil organisms;[1,12] such models explicitly account for different trophic levels or functional groups of biota in the soil.[13–18] Some models that combine an explicit description of the soil biota with a process-based approach have been developed.[19] Food-web models require a detailed knowledge of the Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042721 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter Modeling

biology of the system to be simulated and are usually parameterized for application at specific sites.

FACTORS AFFECTING SOM TURNOVER IN MODELS Rate ‘‘constants’’ (k) are constant for a given set of biotic and abiotic conditions. For nonoptimum environmental circumstances, the simplest way to modify the maximum value of k is by multiplication by a reduction factor m—ranging from 0 to 1. Environmental factors considered by SOM models include temperature, water, pH, nitrogen, oxygen, clay content, cation exchange capacity, type of crop=plant cover, and tillage. Many studies show the effect of temperature on microbially mediated transformations in soil, either expressed as a reduction factor or the Arrhenius equation; but the assumption that SOM decomposition is temperature dependent has been challenged by a study suggesting that old SOM in forest soils does not decompose more rapidly in soils from warmer climates than in soils from colder regions.[20] Recent studies, however, suggest that old SOM is not more resistant than younger pools of SOM.[21,22] Water and oxygen have a major impact on the microbial physiology. Whilst some models simulate O2 concentrations in soil explicitly,[23,24] many define the extent of anaerobiosis based on soil pore space filled with water (WFPS).[25,26] Soil clay content and total SOM are correlated. Various schemes simulate the effect of clay on rate equations to obtain SOM accumulation. Nitrogen is an essential element for microbial growth that will be maximal when enough N is assimilated to maintain the microbial C : N ratio.[27] An overview of the 33 models represented in the GCTE-SOMNET[7,8] including the factors affecting SOM turnover are presented in Table 1.

SOM MODEL EVALUATION There are many reasons for evaluating the performance of a SOM model. Model evaluation shows how well a model can be expected to perform in a given situation, it can help to improve the understanding of the system (especially where the model fails), it can provide confidence in the model’s ability to predict changes in SOM in the future or where there are no data, and it can be used to assess the uncertainties associated with the model’s predictions. Models can be evaluated at a number of different levels. They can be evaluated at the individual process level, at the level of a subset of processes (e.g., net mineralization), or the model’s overall outputs (e.g., changes in total SOM over time) can be tested against measured


1197

laboratory and field data. Models can also be evaluated for their applicability in different situations, e.g., for scaling-up simulated net C storage from a site specific to a regional level.[61] Many examples of different forms of SOM model evaluation are presented elsewhere.[6] In the most comprehensive evaluation of SOM models to date,[61] nine models were tested against 12 data sets from seven long-term experiments representing arable rotations, managed and unmanaged grassland, forest plantations, and natural woodland regeneration. The results showed that six models had significantly lower overall errors [root mean square error (RMSE)] than another group of three models (Fig. 1). The poorer performance of three of the models was related to failures in other parts of the ecosystem models, thus providing erroneous inputs into the SOM module.[61]

SOM MODEL APPLICATION Soil organic matter models are often used as research tools in that they are hypotheses of the dynamics of C and N in soil and can be used to distinguish between competing hypotheses.[49] Another increasing application of SOM models is in agronomy; many SOM models are now being used to improve agronomic efficiency and environmental quality through incorporation into decision support systems, e.g., SUNDIAL-FRS,[56] DSSAT,[35] and APSIM.[29] Soil organic matter models are now used, more than ever, to extrapolate our understanding of SOM dynamics both temporally (in to the future) and spatially (to assess C fluxes from whole regions or continents). An early example of a regional scale application was the use of the CENTURY model to predict the effects of alternative management practices and policies in agroecosystems of central U.S.[63] Since then, many studies have adopted similar methodologies to assess SOM dynamics at the regional,[64] national,[65,66] and global scales.[67–75] Soil organic matter models are increasingly being used by policy makers at the national, regional, or global scales, for example, in the postKyoto debate on the ability of the terrestrial biosphere to store carbon.[76] With such an important role in society, it is important that SOM models are transparent, well evaluated, and well documented. There is still a variety of understanding and different hypotheses incorporated in our current SOM models. Future developments in SOM models will further improve our understanding and allow models to be used truly predictively, without the need for site-specific calibration. These developments will improve estimates of, and reduce, the uncertainty associated with SOM model predictions.

1198

Table 1 Overview of SOM models represented within GCTE-SOMNET in January 2001 Inputs Model

Factors affecting decay rate constants

Meteorology

Soil and plant

Management

ANIMO

Day, week, month

P, AT, Ir, EvW

Des, Lay, Imp, Cl, OM, N, pH

Rot, Ti, Fert, Man, Res, Irr, AtN

T, W, pH, N, O

C, N, W, ST, gas

[28]

APSIM

Day

P, AT, Ir

Lay, W, C, N, BD, Wi, PG, PS

Rot, Ti, Fert, Irr

T, W, pH, N

C, N, W, ST, gas

[29]

Candy

Day

P, AT, Ir

D, Imp, W, N, C, Wi, PD, Nup


T, W, N, Cl

C, N, W, ST, gas

[30]

CENTURY

Month

P, AT

W, Cl, OM, pH, C, N


T, W, N, Cl, pH, Ti

C, BioC, 13C, 14C, N, W, ST, gas

[31]

Chenfang Lin Model

Day

ST

OM, BD, W

Man, Res

T, W, F

C, BioC, gas

[32]

DAISY

Hour, day

P, AT, Ir, EvG

Lay, Cl, C, N, PG, PS


T, W, N, Cl

C, BioC, N, W, ST, gas

[33]

DNDC

Hour, day, month

P, AT

Lay, Cl, OM, pH, BD


T, W, N, Cl, Ti

C, BioC, N, W, ST, gas

[34]

DSSAT

Hour, day, month, year

P, AT, Ir

Des, Lay, Imp, W, Cl, PS, OM, pH, C, N

Rot, Ti, Fert, Man, Res, Irr

T, W, N, Cl, Ti

C, BioC, N, W, ST

[35]

D3R

Day

P, AT

Y, PS

Rot, Ti, Res

T, W, N, Cv, Ti

Decomp. of surface and buried residue

[36]

Ecosys

Minute, hour

P, AT, Ir, WS, RH

Lay, W, Cl, CEC, PS, OM, pH, N, BD, PG, PS


T, W, N, O, Cl, Cv

C, BioC, N, W, ST, pH, Ph, EC, gas, ExCat

[37]

EPIC

Day

P, AT

Lay, Imp, W, Cl, OM, pH, C, BD, Wi

Rot, Ti, Fert, Man, Res, Irr,AtN

T, W, N, pH, Cl, Ce, Cv

C, BioC, N, W, ST

[38]

FERT

Day

P, AT, WS

Des, Lay, W, Cl, OM, pH, C, N, BD, W, Ph, K, Nup, Y, PS

Rot, Ti, Fert, Man, Res, Irr

T, W, N, pH, Cv

C, N, Ph, K

[39]

ForClim-D

Year

P, AT

W, AG

None

T, W

C

[40]

GENDEC

Day, month

ST, W

W, InertC, LQ

Can be used—not essential

T, W, N

C, BioC, N, gas, LQ

[41]

HPM=EFM

Day

P, AT, Ir, WS

W, Cl, PS

Rot, Fert, Irr, AtN

T, W, N

C, BioC, N, W, gas

[42]

ICBM

Day, year

Combination of weather & climate

Many desirable: none essential

C inputs to soil

T, W, Cl

C

[43]

KLIMATSOIL-YIELD

Day, year

P, AT, ST, Ir, EvG, EvS, VPD, SH

Des, Lay, Imp, W, Cl, PS, OM, pH, C, N

Fert, Man, Res, Irr

T, W, N, Cl

C, BioC, N, W, ST

[44]


Soil outputs

References


Timestep

Day

P, AT, Ir

Lay, Inp, W, Cl, CEC, OM, pH, C, N, PS, AS

Fert

T, W, N, pH

C, N, W, ST

[45]

Humus balance

Year

Climate based on P and AT

Des, Lay, PS, OM, pH, C, N

Rot, Fert, Man

N, H, Cl, Cv

C, N

[46]

MOTOR

User specified

P, AT, EvG

Des, OM

Rot, Ti, Fert, Man

T, W, N, Cl, Ti

C, BioC, 13C, 14C, gas

[47]

NAM SOM

Year

P, AT

Des, PS, OM, Ero

Man, Res

T, W, Cl, Cv

C, BioC

[48]

NCSOIL

Day

ST (P, AT)

W, OM, C, N

Fert, Man, Res

T, W, N, pH, Cl, Ti

C, BioC, 14C, N, 15N, gas

[49]

NICCE

Hour, day

P, AT, Ir, WS

Imp, OM, C, N, W, TC, PG

Fert, Man, Res, Irr, AtN

T, W, Cl, N

C, BioC, 13C, 14C, N, 15N, W, ST, gas

[50]

O’Brien model

Year

None

Lay, C, 14C

None

None

C, 14C

[51]

O’Leary model

Day

P, AT

Lay, W, Cl, pH, N

Ti, Fert, Res

T, W, N, Cl, Ti

C, BioC, N, W, ST, gas, ResC, ResN

[52]

Q-soil

Year

Optional

C, N

Rot, Fert, Man, Res, AtN

T, W, N

C, BioC, 13C, N

[10]

RothC

Month

P, AT, EvW

Cl, C, InertC (can be estimated)

Man, Res, Irr

T, W, Cl, Cv

C, BioC, gas, 14C

[53]

SOCRATES

Week

P, AT

CEC, Y

Rot, Fert, Res

T, W, N, Cv, Ce

C, BioC, gas

[54]

SOMM

Day

P, ST

OM, N, AshL, NL

Man

T, W, N

C, N, gas

[55]

Sundial

Week

P, AT, EvG

Imp, Cl, W, Y

Rot, Fert, Man, Res, Irr, AtN

T, W, N, Cl

C, BioC, N, 15N, W, gas

[56]

Verberne

Day

P, AT, Ir, WS, EvS

Des, W, Cl, PS, OM, C, N

Man, AtN

T, W, N, Cl

C, BioC, N, W

[57]

VOYONS

Day, week, month

P, ST

Cl, OM, C, N

Fert, Man, Res, Irr, AtN

T, W, Cl

C, BioC, 13C, 14C, N, gas

[58]

Wave

Day

P, AT, Ir, EvG

Lay, OM, C, N, W, PG


T, W, N

C, N, W, ST, gas

[59]


CNSP pasture model

P, precipitation; AT, air temperature; ST, soil temperature; Ir, irradiation; EvW, evaporation over water; EvG, evaporation over grass; EvS, evaporation over bare soil; WS, wind speed; RH, relative humidity; VPD, vapor pressure deficit; SH, sun hours; Des, soil description; Lay, soil layers; Imp, depth of impermeable layer; Cl, clay content; OM, organic matter content; N, soil nitrogen content= dynamics; C, soil carbon content= dymanics; InertC, soil inert carbon content; W, Soil water characteristics; Wi, wilting point; PD, soil particle size distribution; CEC, cation exchange capacity; Ero, annual erosion losses; BD, soil bulk density; TC, thermal conductivity; PG, plant growth characteristics; PS, plant species composition; AS, animal species present; AG, animal growth characteristics; Y, yield; Nup, plant nitrogen uptake; LQ, litter quality; AshL, ash content of litter; NL, N content of litter; Rot, rotation; Ti, tillage practice; Fert, inorganic fertilizer applications; Man, organic manure applications; Res, residue management; Irr, irrigation; AtN, atmospheric nitrogen inputs; T, temperature; W, water; N, nitrogen; O, oxygen; Cl, clay; Ce, cation exchange capacity; Cv, cover crop; Ti, tillage; F, Fauna; BioC, Biomass carbon; 13C, 13C dynamics; 14C, 14C dynamics; 15N, 15N dynamics; gas, gaseous losses (e.g., CO2, N2O, and N2); ResC, surface residue carbon; ResN, surface residue nitrogen; Ph, phosphorus dynamics; K, potassium dynamics; EC, electrical conductivity; ExCat, exchangeable cations. NB, N in the soil inputs and outputs section is used to denote all aspects of the N cycle. Further details regarding optimum decay conditions, SOM components, rate constants, methods of pool fitting, and refractory SOM are given elsewhere. (Refs.[6,60] and a metadatabase of all models is available Ref.[7].) 1199


1200

Fig. 1 Overall root mean square error value for nine SOM models when simulating changes in total soil organic carbon in up to 12 datasets from seven long term experiments. The RMSE values of the models with the same letter (a or b) do not differ significantly (two sample, two tailed t-test; p > 0.05), but the RMSE values of the two groups (a and b) do differ significantly (two sample, two tailed t-test; p < 0.05).[62]

CONCLUSIONS Soil organic matter models are used for many purposes in soil science, including use as tools to synthesize and explore, explain, and extrapolate experimental data, use for making projections of SOM behavior under current and future environmental conditions, and use for supporting decision making at many levels, by many users. Soil organic matter models remain one of our most important tools for improving our understanding soil of organic matter dynamics.

REFERENCES 1. Paustian, K. Modelling soil biology and biochemical processes for sustainable agricultural research. In Soil Biota. Management in Sustainable Farming Systems; Pankhurst, C.E., Doube, B.M., Gupta, V.V.S.R., Grace, P.R., Eds.; CSIRO Information Services: Melbourne, 1994; 182–193. 2. Jenny, H. Factors of Soil Formation. A System of Quantitative Pedology; McGraw-Hill: New York, 1941. 3. Beek, J.; Frissel, M.J. Simulation of Nitrogen Behaviour in Soils; Pudoc: Wageningen, The Netherlands, 1973. 4. Jenkinson, D.S. Studies on the decomposition of plant material in soil. V. J. Soil Sci. 1977, 28, 424–434. 5. McGill, W.B. Review and classification of ten soil organic matter (SOM) models. In Evaluation of Soil Organic Matter Models Using Existing, Long-Term Datasets; NATO ASI I38; Smith, J.U., Powlson, D.S., Smith, P., Eds.; Springer-Verlag: Berlin, 1996; 111–133.



6. Molina, J.A.E.; Smith, P. Modeling carbon and nitrogen processes in soils. Adv. Agron. 1998, 62, 253–298. 7. Global Change and Terrestrial Ecosystems (GCTE) Soil Organic Matter Network (SOMNET) database. Available at: http:==www.iacr.bbsrc.ac.uk=aen=somnet= index.htm. 8. Smith, P.; Smith, J.U.; Powlson, D.S., Eds. Soil Organic Matter Network (SOMNET): 1996 Model and Experimental Metadata; GCTE Report 7; GCTE Focus 3 Office: Wallingford, U.K., 1996; 255 pp. 9. Smith, P.; Powlson, D.S.; Smith, J.U.; Glendining, M.J. The GCTE SOMNET: a global network and database of soil organic matter models and long-term datasets. Soil Use Manage. 1996, 108, 57. ˚ gren, G.I. Theoretical analyses of the 10. Bosatta, E.; A interactions between inorganic nitrogen and soil organic matter. Eur. J. Soil Sci. 1995, 76, 109–114. ˚ gren, G.I. Theoretical analysis of micro11. Bosatta, E.; A bial biomass dynamics in soils. Soil Biol. Biochem. 1994, 26, 143–148. 12. Smith, P.; Andreń, O.; Brussaard, L.; Dangerfield, M.; Ekschmitt, K.; Lavelle, P.; Tate, K. Soil biota and global change at the ecosystem level: describing soil biota in mathematical models. Global Change Biol. 1998, 4, 773–784. 13. Hunt, H.W.; Coleman, D.C.; Cole, C.V.; Ingham, R.E.; Elliott, E.T.; Woods, L.E. Simulation model of a food web with bacteria, amoebae, and nematodes in soil. In Current Perspectives in Microbial Ecology; Klug, M.J., Reddy, C.A., Eds.; American Society for Microbiology: Washington, DC, 1984; 346–352. 14. Hunt, H.W.; Coleman, D.C.; Ingham, E.R.; Ingham, R.E.; Elliot, E.T.; Moore, J.C.; Rose, S.L.; Reid, C.P.; Morley, C.R. The detrital food web in a shortgrass prairie. Biol. Fertility Soils 1987, 3, 57–68. 15. Hunt, H.W.; Trlica, M.J.; Redente, E.F.; Moore, J.C.; Detling, J.K.; Kittel, T.G.F.; Wlater, D.E.; Fowler, M.C.; Klein, D.A.; Elliot, E.T. Simulation model for the effects of climate change on temperate grassland ecosystems. Ecol. Model. 1991, 53, 205–246. 16. de Ruiter, P.C.; Van Faassen, H.G. A comparison between an organic matter dynamics model and a food web model simulating nitrogen mineralization in agroecosystems. Eur. J. Agron. 1994, 3, 347–354. 17. de Ruiter, P.C.; Van Veen, J.A.; Moore, J.C.; Brussaard, L.; Hunt, H.W. Calculation of nitrogen mineralization in soil food webs. Plant Soil 1993, 157, 263–273. 18. de Ruiter, P.C.; Neutel, A.-M.; Moore, J.C. Energetics and stability in belowground food webs. In Food Webs, Integration of Patterns and Dynamic; Polis, G.A., Winemiller, K.O., Eds.; Chapman & Hall: New York, 1995; 201–210. 19. McGill, W.B.; Hunt, H.W.; Woodmansee, R.G.; Reuss, J.O. PHOENIX, a model of the dynamics of carbon and nitrogen in grassland soil. In Terrestrial Nitrogen Cycles. Processes, Ecosystem Strategies and Management Impacts; Clark, F.E., Roswall, T., Eds.; Ecological Bulletins 1981; Vol. 33, 49–115. 20. Giardina, C.P.; Ryan, M.G. Evidence that decomposition rates of organic carbon in mineral soil do not vary with temperature. Nature 2000, 393, 249–252.


21. Fang, C.; Smith, P.; Moncrieff, J.; Smith, J.U. Similar response of labile and resistant soil organic matter pools to changes in temperature. Nature 2005, 433, 57–59. 22. Knorr, W.; Prentice, I.C.; House, J.I.; Holland, E.A. Long-term sensitivity of soil carbon turnover to warming. Nature 2005, 433, 298–301. 23. Grant, R.F. A technique for estimating denitrification rates at different soil temperatures, water contents and nitrate concentrations. Soil Sci. 1991, 152, 41–52. 24. Sierra, J.; Renault, P. Respiratory activity and oxygen distribution in natural aggregates in relation to anaerobiosis. Soil Sci. Soc. Am. J. 1996, 60, 1428–1438. 25. Skopp, J.; Jawson, M.D.; Doran, J.W. Steady-state aerobic microbial activity as a function of soil water content. Soil Sci. Soc. Am. J. 1990, 54, 1619–1625. 26. Doran, J.W.; Mielke, L.N.; Stamatiadis, S. Microbial activity and N cycling as regulated by soil water-filled pore space. Proceedings of the 11th Conference International Soil Tillage Research Organization, Edinburg, Scotland, 1988; Vol. 1, 49–54. 27. Molina, J.A.E.; Clapp, C.E.; Shaffer, M.J.; Chichester, F.W.; Larson, W.E. NCSOIL, a model of nitrogen and carbon transformations in soil: description, calibration, and behavior. Soil Sci. Soc. Am. J. 1983, 47, 85–91. 28. Rijtema, P.E.; Kroes, J.G. Some results of nitrogen simulations with the model ANIMO. Fert. Res. 1991, 27, 189–198. 29. McCown, R.L.; Hammer, G.L.; Hargreaves, J.N.G.; Holzworth, D.P.; Freebairn, D.M. APSIM: a novel software system for model development, model testing and simulation in agricultural systems research. Agr. Syst. 1996, 50, 255–271. 30. Franko, U. Modelling approaches of soil organic matter turnover within the CANDY system. In Evaluation of Soil Organic Matter Models Using Existing, Long-Term Datasets; NATO ASI I38; Powlson, D.S., Smith, P., Smith, J.U., Eds.; Springer-Verlag: Berlin, 1996; 247–254. 31. Parton, W.J.; Stewart, J.W.B.; Cole, C.V. Dynamics of C, N, P, and S in grassland soils: a model. Biogeochemistry 1987, 5, 109–131. 32. Lin, C.; Liu, T.S.; Hu, T.L. Assembling a model for organic residue transformation in soils. Proc. Natl. Council (Taiwan) B 1987, 11, 175–186. 33. Mueller, T.; Jensen, L.S.; Hansen, S.; Nielsen, N.E. Simulating soil carbon and nitrogen dynamics with the soil–plant–atmosphere system model DAISY. In Evaluation of Soil Organic Matter Models Using Existing, Long-Term Datasets; NATO ASI I38; Powlson, D.S., Smith, P., Smith, J.U., Eds.; Springer-Verlag: Berlin, 1996; 275–281. 34. Li, C.; Frolking, S.; Harriss, R. Modelling carbon biogeochemistry in agricultural soils. Global Biogeochem. Cycles 1994, 8, 237–254. 35. Hoogenboom, G.; Jones, J.W.; Hunt, L.A.; Thornton, P.K.; Tsuji, G.Y. An Integrated Decision Support System for Crop Model Applications, Paper 94-3025, ASAE Meeting, Missouri, June 1994; 23 pp. 36. Douglas, C.L., Jr.; Rickman, R.W. Estimating crop residue decomposition from air temperature, initial


1201

37.

38.

39.

40.

41.

42. 43.

44.

45.

46.

47.

48.

49.

50.

51.

52.

53.

nitrogen content, and residue placement. Soil Sci. Soc. Am. J. 1992, 56, 272–278. Grant, R.F. Dynamics of energy, water, carbon and nitrogen in agricultural ecosystems: simulation and experimental validation. Ecol. Model. 1995, 81, 169–181. Williams, J.R. The erosion-productivity impact calculator (EPIC) model: a case history. Phil. Trans. Roy. Soc. Lond. B 1990, 329, 421–428. Kan, N.A.; Kan, E.E. Simulation model of soil fertility. Physiol. Biochem. Cultivated Plants 1991, 23, 3–16, (in Russian). Perruchoud, D.O. Modeling the Dynamic of Non-living Organic Carbon in a Changing Climate: A Case Study for Temperate Forests. Ph.D., Thesis, ETH Diss. No. 11900, 1996; 196 pp. Moorhead, D.L.; Reynolds, J.F. A general model of litter decomposition in the northern Chihuahuan desert. Ecol. Model. 1991, 56, 197–219. Thornley, J.H.M.; Verberne, E.L.J. A model of nitrogen flows in grassland. Plant Cell Environ. 1989, 12, 863–886. Andreń, O.; Ka¨tterer, T. ICBM—the introductory carbon balance model for exploration of soil carbon balances. Ecol. Appl. 1997, 7, 1226–1236. Sirotenko, O.D. The USSR climate–soil–yield simulation system. Meteorologia i Gidrologia 1991, 4, 67–73, (in Russian). McCaskill, M.; Blair, G.J. A model of S, P and N uptake by a perennial pasture. I. Model construction. Fert. Res. 1990, 22, 161–172. Schevtsova, L.K.; Mikhailov, B.G. Control of Soil Humus Balance Based on Statistical Analysis of Long-Term Field Experiments Database; VIUA: Moscow, 1992, (in Russian). Whitmore, A.P.; Klein-Gunnewiek, H.; Crocker, G.J.; Klı´r, J.; Ko¨rschens, M.; Poulton, P.R. Simulating trends in soil organic carbon in long-term experiments using the Verberne=MOTOR model. Geoderma 1997, 81, 137–151. Ryzhova, I.M. Analysis of sensitivity of soil-vegetation systems to variations in carbon turnover parameters based on a mathematical model. Eurasian Soil Sci. 1993, 25, 43–50. Molina, J.A.E.; Hadas, A.; Clapp, C.E. Computer simulation of nitrogen turnover in soil and priming effect. Soil Biol. Biochem. 1990, 22, 349–353. Van Dam, D.; Van Breemen, N. NICCE—a model for cycling of nitrogen and carbon isotopes in coniferous forest ecosystems. Ecol. Model. 1995, 79, 255–275. O’Brien, B.J. Soil organic carbon fluxes and turnover rates estimated from radiocarbon measurements. Soil Biol. Biochem. 1984, 16, 115–120. O’Leary, G.J. Soil Water and Nitrogen Dynamics of Dryland Wheat in the Victorian Wimmera and Malle. Ph.D., Thesis, University of Melbourne, 1994; 332 pp. Coleman, K.; Jenkinson, D.S.; Crocker, G.J.; Grace, P.R.; Klı´r, J.; Ko¨rschens, M.; Poulton, P.R.; Richter, D.D. Simulating trends in soil organic carbon in longterm experiments using RothC-23.6. Geoderma 1997, 81, 29–44.

1202

54. Grace, P.R.; Ladd, J.N. SOCRATES v2.00 User Manual; Co-operative Research Centre for Soil and Land Management: Glen Osmond, South Australia, 1995. 55. Chertov, O.G.; Komarov, A.S. SOMM—a model of soil organic matter and nitrogen dynamics in terrestrial ecosystems. In Evaluation of Soil Organic Matter Models Using Existing, Long-Term Datasets; Powlson, D.S., Smith, P., Smith, J.U., Eds.; NATO ASI I38; SpringerVerlag: Berlin, 1996; 231–236. 56. Smith, J.U.; Bradbury, M.J.; Addiscott, T.M. SUNDIAL: simulation of nitrogen dynamics in arable land. A user-friendly, PC-based version of the Rothamsted nitrogen turnover model. Agron. J. 1996, 88, 38–43. 57. Verberne, E.L.J.; Hassink, J.; de Willigen, P.; Groot, J.R.R.; van Veen, J.A. Modelling soil organic matter dynamics in different soils. Neth. J. Agric. Sci. 1990, 38, 221–238. 58. Andre´, M.; Thiery, J.M.; Courmac, L. ECOSIMP model: prediction of CO2 concentration changes and carbon status in closed ecosystems. Adv. Space Res. 1992, 14, 323–326. 59. Vanclooster, M.; Viaene, P.; Diels, J.; Feyen, J. A deterministic evaluation analysis applied to an integrated soil-crop model. Ecol. Model. 1995, 81, 183–195. 60. Falloon, P.; Smith, P. Modelling refractory organic matter—a review. Biol. Fert. Soils 2000, 30, 388–398. 61. Izaurralde, R.C.; Haugen-Kozyra, K.H.; Jans, D.C.; McGill, W.B.; Grant, R.F.; Hiley, J.C. Soil organic carbon dynamics: measurement, simulation and site to region scale-up. In Assessment Methods for Soil Carbon. Advances in Soil Science; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, FL, 2001; 553–575. 62. Smith, P.; Smith, J.U.; Powlson, D.S.; McGill, W.B.; Arah, J.R.M.; Chertov, O.G.; Coleman, K.; Franko, U.; Frolking, S.; Jenkinson, D.S.; Jensen, L.S.; Kelly, R.H.; Klein-Gunnewiek, H.; Komarov, A.; Li, C.; Molina, J.A.E.; Mueller, T.; Parton, W.J.; Thornley, J.H.M.; Whitmore, A.P. A comparison of the performance of nine soil organic matter models using datasets from seven long-term experiments. Geoderma 1997, 81, 153–225. 63. Donigian, A.S., Jr.; Barnwell, T.O., Jr.; Jackson, R.B., IV; Patwardhan, A.S.; Weinrich, K.B.; Rowell, A.L.; Chinnaswamy, R.V.; Cole, C.V. Assessment of Alternative Management Practices and Policies Affecting Soil Carbon in Agroecosystems of the Central United States, US EPA Report EPA=600=R-94=067; Athens, 1994; 194 pp. 64. Falloon, P.; Smith, P.; Smith, J.U.; Szabo´, J.; Coleman, K.; Marshall, S. Regional estimates of carbon sequestration potential: linking the Rothamsted carbon model to GIS databases. Biol. Fert. Soil 1998, 27, 236–241.



65. Lee, J.J.; Phillips, D.L.; Liu, R. The effect of trends in tillage practices on erosion and carbon content of soils in the US corn belt. Water Air Soil Poll. 1993, 70, 389–401. 66. Parshotam, A.; Tate, K.R.; Giltrap, D.J. Potential effects of climate and land-use change on soil carbon and CO2 emissions from New Zealand’s indigenous forests and unimproved grasslands. Weather Climate 1996, 15, 3–12. 67. Post, W.M.; Emanuel, W.R.; Zinke, P.J.; Stangenberger, A.G. Soil carbon pools and world life zones. Nature 1982, 298, 156–159. 68. Post, W.M.; Pastor, J.; Zinke, P.J.; Staggenberger, A.G. Global patterns of soil nitrogen storage. Nature 1985, 317, 613–616. 69. Post, W.M.; King, A.W.; Wullschleger, S.D. Soil organic matter models and global estimates of soil organic carbon. In Evaluation of Soil Organic Matter Models Using Existing, Long-Term Datasets; Powlson, D.S., Smith, P., Smith, J.U., Eds.; NATO ASI I38; Springer-Verlag: Berlin, 1996; 201–222. 70. Potter, C.S.; Randerson, J.T.; Field, C.B.; Matson, P.A.; Vitousek, P.M.; Mooney, H.A.; Klooster, S.A. Terrestrial ecosystem production: a process model based on satellite and surface data. Global Biogeochem. Cycles 1993, 7, 811–841. 71. Schimel, D.S.; Braswell, B.H., Jr.; Holland, E.A.; McKeown, R.; Ojima, D.S.; Painter, T.H.; Parton, W.J.; Townsend, J.R. Climatic, edaphic, and biotic controls over storage and turnover of carbon in soils. Global Biogeochem. Cycles 1994, 8, 279–293. 72. Goto, N.; Sakoda, A.; Suzuki, M. Modelling soil carbon dynamics as a part of the carbon cycle in terrestrial ecosystems. Ecol. Model. 1993, 74, 183–204. 73. Esser, G. Modelling global terrestrial sources and sinks of CO2 with special reference to soil organic matter. In Soils and the Greenhouse Effect; Bouwman, A.F., Ed.; John Wiley & Sons: NewYork, 1990; 247–261. 74. Goldewijk, K.K.; van Minnen, J.G.; Kreileman, G.J.J.; Vloedbeld, M.; Leemans, R. Simulating the carbon flux between the terrestrial environment and the atmosphere. Water Air Soil Poll. 1994, 76, 199–230. 75. Melillo, J.M.; Kicklighter, D.W.; McGuire, A.D.; Peterjon, W.T.; Newkirk, K.M. Global change and its effect on soil organic carbon stocks. In Role of Nonliving Organic Matter in the Earth’s Carbon Cycle; Zepp, R.G., Sonntag, C.H., Eds.; John Wiley & Sons: New York, 1995; 175–189. 76. IPCC. Land Use, Land-Use Change, and Forestry. A Special Report of the IPCC; Cambridge University Press: Cambridge, U.K., 2000; 377 pp.

Organic Matter Structure and Characterization Georg Guggenberger Institute of Soil Science and Soil Geography, University of Bayreuth, Bayreuth, Germany

INTRODUCTION Soil organic matter (SOM) encompasses all biologicallyderived organic material found in the soil or on its surface irrespective of 1) source; 2) whether it is living or dead; or 3) stage of decomposition, but excluding the aboveground portion of living plants.[1] This implies large structural heterogeneity and close linkage to SOM functions. Litter composition and its changes during biotic and abiotic transformation are key variables in the processes and the size of carbon sequestration in soil. In the context of the Kyoto Protocol, analysis of SOM structure helps us to understand these processes and to predict changes in the sign and magnitude of terrestrial carbon fluxes in a changing environment. To reduce SOM heterogeneity, different components of SOM need to be separated into entities that differ in terms of source, composition, and turnover. Recent evidence suggests that the classical chemical fractionation into fulvic acids, humic acids, and humin according to solubility characteristics in dilute acid and base (for standard procedure see Ref.[2]) is not useful in this respect.[3] Fractionations that are more promising rely mostly on physical fractionations according to particle size or density,[4,5] and the analysis of dissolved organic matter in the soil solution.[6] ANALYTICAL METHODS The analytical approaches can be divided into degradative methods involving chemolysis, thermolysis, or thermochemolysis, and noninvasive spectroscopic techniques such as nuclear magnetic resonance (NMR) spectroscopy (Table 1). Degradative methods provide molecular-level information on specific organic compounds accessible to the degradative step whereas spectroscopic techniques inform on the bulk composition of SOM. Since all analytical approaches have their drawbacks, a comprehensive picture of the SOM structure can only be obtained using a combination of various methods.[7] Degradative Methods Chemolysis of SOM involves various degradation procedures (hydrolysis, oxidation, extraction) that are Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001588 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

selective in their attack on specific molecular structures (Table 1). The techniques are well suited to follow the often subtle changes in the composition of plant- and microbial-derived biopolymers during microbial transformation processes, in particular when different soil fractions are comparatively investigated.[8,9] But only a part of SOM present in specific structures—as determined by NMR spectroscopy—can be identified with these techniques, and information on the original macromolecular structure of SOM is limited.[10] Analytical pyrolysis uses thermal degradation to cleave bonds in the organic macromolecules and enables a sensitive and rapid characterization of organic constituents.[11] Problems may arise due to the complicated pyrolysis behavior of many organic compounds, in particular in the presence of catalytic minerals. Thermal secondary reaction causes considerable modification of the organic compound, which may bias the interpretation of the pyrolysis products with respect to their mother compounds.[12,13] Some of the problems can be solved by thermochemolysis utilizing tetramethylammonium hydroxide (TMAH).[13,14] Hydroxyl and carboxyl groups are converted into their respective methyl ethers and methyl esters to avoid fragmentation of aliphatic and benzenecarboxylic acids. Spectroscopic Techniques Nuclear magnetic resonance spectroscopy has a large potential to analyze the different chemical species of 1 H, 13C, 15N, and 31P nuclei in solution and solid (not 1H) state. With the cross-polarization magic angle spinning (CP MAS) pulse sequence, the chemical structure of C and N can be characterized in situ nondestructively, and within a feasible time of analysis.[15,16] However, even with completion of detailed experiments on rates of signal generation and relaxation for each type of carbon,[17] quantitative analysis of NMR data is not possible in the presence of paramagnetic species. Samples containing charred materials cannot be quantified by the CP technique; instead, the C nuclei must be directly excited using a Bloch decay sequence.[18] Additional information on the structure of soil organic carbon (SOC) can be obtained by specific pulse sequences such as interrupted decoupling (ID), proton spin relaxation editing (PSRE), and mixing of proton spins (MOPS) (for 1203

1204

Table 1 Important degradative and nondegradative techniques to study SOM structure Technique Chemolytic techniques

Acid hydrolysis

Components addressed Address defined biomolecules released mostly as their monomer units from the organic matrix by an appropriate chemical treatment Polysaccharides

Proteinaceous compounds

Amino sugars

Solvent extraction

Extractable lipids

Saponification

Cutin and suberin components

CuO oxidation

Lignin

HNO3 oxidation

Analytical pyrolysis


Pyrogenic carbon

Products of SOM released by thermal degradation

Analysis of defined organic components of plant, microbial and pyrogenic origin Often well-suited as biomarkers Distinguishes different types of polysaccharides by sequential extraction Pattern of monosaccharides released gives information on the source (plant vs. microorganisms) Comprises the dominant form of organic nitrogen in soil Assessment of D=L ratios Traces microbial input to SOM Distinguishes between fungal and bacterial residues Comprises hydrophobic aliphatic compounds of various origins Detects aliphatic biopolymers derived from vascular plants Traces plant input to SOM Informs about type of plant Informs about lignin decomposition in soil Detects aliphatic biopolymers derived from vascular plants Traces the highly aromatic core of charred organic components Pattern of pyrolysis products allows detailed information on the chemical composition of the parent molecules Well suited for biomarker analysis

Drawbacks Only about 60% (fresh organic residues) to about 20% (transformed, mineral-associated compounds) of SOM is characterized Not all polysaccharides are hydrolyzable

References [7,34]

[8,35]

Yield differs for different types of with organic matter

Only a part of the proteinaceous compounds is hydrolyzable Source of some amino sugars is not well defined

[36,37]

Complex and polymerized lipids are only scarcely accessible Does not comprise nonester plant biomacromolecules Only cleaves arylether bondings Yield is not defined and decreases with increasing lignin decomposition Does not comprise nonester plant biomacromolecules Yield varies with the type of the charred organic components

[39,40]

Volatilization of different types of pyrolysis fragments varies Pyrolysis fragments may have different sources Secondary pyrolysis reactions may occur

[22,38]

[25] [41,42]

[43] [44,45]

[11,12]

Organic Matter Structure and Characterization

Cutin and suberin components

Applications

Py-FIMS

Pyrolysis products that are volatile to be separated by gas chromatography Pyrolysis products are treated by soft ionization technique

Solution 13C, 1H, 15N, and 31P NMR spectroscopy

In particular aliphatic and benzenecarboxylic acids as their methyl esters Analysis of all C, H, N, or P species within a solid or soluble sample Analysis of C, H, N, and P species in aqueous samples

CP MAS 13C, and N spectroscopy

Analysis of C and N species in solid samples

Thermochemolysis with TMAH NMR spectroscopy

15

BD

13

C NMR spectroscopy

IR spectroscopy (FTIR and DRIFT)

Analysis of C species in solid samples Analysis of C species in solid samples

Separates pyrolysis products that have the same mass=charge (m=z) ratio Soft ionization produces predominantly molecular ions of the pyrolysis products Pyrolysis=methylation renders polar products volatile for gas chromatographic separation Signal intensity relates to the concentration of nuclei creating the signal Composition of C, H, and N species in the soil solution and in alkaline extracts (humic and fulvic acids) Characterization of the chemical structure of organic C and N in situ and nondestructively

Direct excitation of 13C spins enables analysis of charred materials Informs on the type of atoms to which C is bonded and on the nature of the bond

Nonvolatile pyrolysis products cannot be detected

[13]

Volatilization decreases with increasing polarity of the fragments Reproducibility needs to be improved

[11]

Most applications are not quantitative Insoluble compounds are not detected

Limited quantitative interpretation caused by paramagnetic species, by different C and N relaxation, by spinning side bands, and by the cross-polarization technique Very time consuming

Reveals only a few well-resolved peaks Signals from minerals often dominate

[46,47]

[15–17,48]

[49]


Py-GC MS

[15,16]

[18]

[50] [19,50]

Key: SOM ¼ soil organic matter, Py-GC MS ¼ pyrolysis-gas chromatography mass spectrometry, Py-FIMS ¼ pyrolysis-field ionization mass spectrometry, TMAH ¼ tetramethylammonium hydroxide, NMR spectroscopy ¼ nuclear magnetic resonance spectroscopy, CP MAS NMR spectroscopy ¼ cross-polarization magic angle spinning nuclear magnetic resonance spectroscopy, BD NMR spectroscopy ¼ bloch decay nuclear magnetic resonance spectroscopy, IR spectroscopy ¼ infrared spectroscopy, FTIR ¼ fourier transform infrared spectroscopy, DRIFT ¼ diffuse reflectance infrared fourier transform spectroscopy.

1205


1206

literature see Ref.[1].) Of the other spectroscopic techniques, infrared spectroscopy provides information on the type of atoms to which C is bonded and on the nature of the bond,[19] whereas electron spin resonance spectroscopy gives information about free radicals in SOM.[20]

CHEMICAL STRUCTURE OF SOIL ORGANIC MATTER Individual Plant and Microbial Components Individual components derived from plants (primary resources) and microorganisms (secondary resources) can be considered as the parent material of SOM. Within plant and microbial tissues, polysaccharides are generally the most important organic components. Analysis of sugar monomers released by hydrolysis or pyrolysis revealed that crystalline and noncrystalline


plant polysaccharides are rapidly decomposed and microbial polysaccharides accumulate in soil.[8,21] Most amino sugars in soil are also of microbial origin, predominantly from fungal chitin and bacterial peptidoglycan.[22] An important biopolymer in soil is vascular plant-derived lignin with its p-hydroxyphenyl, guaiacyl, and syringyl monomeric units. The pattern of lignin monomers can be used to identify the primary resource of SOM; in contrast to angiosperm lignin, gymnosperm lignin does not contain syringyl units. In soil, lignin is decomposed in aerobic environments via side-chain oxidation and ring opening.[23,24] This results in a shortening of alkyl side-chains and an increase in carbonyl and carboxyl groups. Thus, parts of the lignin degradation products are water soluble[23] and represent, together with the polysaccharides, the majority of the dissolved organic matter in soil.[6] The third major component is free and bound lipids of plant and microbial origin, including waxes, organic acids, steroids, glycerides, and phospholipids.[3,25]

Fig. 1 (A) Thermograms for compound classes evolved by pyrolysis-field ionization from clay (above), fine silt (center), and medium silt (below) from a humus formation experiment after 13 and 34 years of soil development (From Ref.[11]) and (B) solid-state 13C NMR spectra of the clay fraction isolated from soils of different pedogenesis. (From Ref.[10].)



Plant cutin and suberin (insoluble polyesters) can be easily decomposed,[26] while nonsaponifiable cutan and suberan (insoluble nonpolyesters) appear to be resistant to degradation.[27] Some aliphatic compounds of microbial origin are also selectively preserved.[28]

1207

and concepts on SOM structure, the excellent reviews of Ko¨gel-Knabner,[10] Baldock and Nelson,[1] and Hedges et al.[32] are recommended.

REFERENCES The Soil Organic Matter Continuum Soil organic matter is generally composed of the above-mentioned plant and microbial residues, and their transformation products.[7,29] The latter refers largely to the term ‘‘humic substances’’ as frequently used in older references e.g.,[30,31] Recently, there has been a shift in paradigm away from humic substances. Organic matter degradation is considered to be predominantly a process of attrition, during which relatively resistant biomolecules are selectively concentrated.[32] In surface soil horizons with large inputs of plant residues, the SOM continuum varying from fresh residues to highly degraded components is dominated by the former material.[7] In contrast, SOM in subsoil horizons is enriched by highly-altered, recalcitrant organic materials that can only be characterized to a minor extent by chemolytic treatments.[32] Accessibility of these compounds is limited to (chemo)thermolytic and spectroscopic methods. In aging SOM, Schulten, Leinweber, and Reuter[33] and Schulten and Leinweber[11] observed that increasing amounts of thermal energy applied in pyrolysis-field ionization mass spectrometry were required for volatilizing lignin dimers, alkylaromatics, and lipids (Fig. 1A). The authors suggested that the development of stronger chemical bonds was due to either formation of threedimensional crosslinking by aryl–alkyl combinations or by formation of strong organic–mineral complexes. Ko¨gel-Knabner[7] concluded from ID pulse sequences at 13C NMR spectroscopy that cross-linked aliphatic compounds accumulate during SOM decomposition. The pedogenic environment has a large impact on the structure of SOM. This can be particularly confined by solid-state 13C NMR spectroscopy of mineralassociated organic matter, while spectra obtained on bulk samples are influenced by plant residues that are rather similar in composition.[10] The Mollisol in Fig. 1B is characterized by about similar contributions of alkyl (0–50 ppm), O-alkyl (50–110 ppm), aromatic (110–160 ppm) and carboxyl/amide (160–200 ppm) carbon. In contrast, Alfisols and Ultisols show a high proportion of alkyl and O-alkyl C, and Spodosols are dominated by alkyl C. Future work may benefit from the application of microscopic and nondestructive microspectroscopic techniques to investigate SOM within its mineral and microbiological soil environment. For in-depth information on modern methodological approaches


1. Baldock, J.A.; Nelson, P.N. Soil organic matter. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, FL, 2000; B-25–B-84. 2. Swift, R.S. Organic matter characterisation. In Methods of Soil Analysis. Part 3. Chemical Methods; Sparks, D.L., Ed.; Soil Science Society of America: Madison, WI, 1996; 1011–1069. 3. Stevenson, F.J.; Elliott, E.T. Methodologies for assessing the quantity and quality of soil organic matter. In Dynamics of Soil Organic Matter in Tropical Ecosystems; Coleman, D.C., Oades, J.M., Uehara, G., Eds.; University of Hawaii Press: Honolulu, HI, 1989; 173–199. 4. Christensen, B.T. Carbon in primary and secondary organomineral complexes. In Structure and Organic Matter Storage in Agricultural Soils; Carter, M.R., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1996; 97–165. 5. Golchin, A.; Oades, J.M.; Skjemstad, J.O.; Clarke, P. Soil structure and carbon cycling. Aus. J. Soil Res. 1994, 32, 1043–1068. 6. Guggenberger, G.; Zech, W.; Schulten, H.-R. Formation and mobilization pathways of dissolved organic carbon: evidence from chemical structural studies of organic carbon fractions in acid forest floor solutions. Org. Geochem. 1994, 21, 51–66. 7. Ko¨gel-Knabner, I. Biodegradation and humification processes in forest soils. In Soil Biochemistry; Bollag, J.-M., Stotzky, G., Eds.; Marcel Dekker: New York, NY, 1993; Vol. 8, 101–137. 8. Guggenberger, G.; Christensen, B.T.; Zech, W. Land-use effects on the composition of organic matter in particle-size separates of soils: I. lignin and carbohydrate signature. Eur. J. Soil Sci. 1994, 45, 449–458. 9. Hedges, J.I.; Oades, J.M. Comparative organic geochemistries of soils and sediments. Org. Geochem. 1997, 27, 319–361. 10. Ko¨gel-Knabner, I. Analytical approaches for characterizing soil organic matter. Org. Geochem. 2000, 31, 609–625. 11. Schulten, H.-R.; Leinweber, P. Characterization of humic and soil particles by analytical pyrolysis and computer modeling. J. Anal. Appl. Pyrol. 1996, 38, 1–53. 12. Saiz-Jimenez, C. Analytical pyrolysis of humic substances: pitfalls, limitations and possible solutions. Environ. Sci. Technol. 1994, 28, 1773–1780. 13. Van Bergen, P.; Flannery, M.B.; Poulton, P.R.; Evershed, R.P. Organic geochemical studies of soils from Rothamsted experimental station: III. Nitrogencontaining organic matter in soil from geescroft wilderness. In Fate of N-Containing Macromolecules in the Biosphere and Geosphere; Stankiewicz, B.A., van Bergen, P.F., Eds.; Symposium Series 707; American

1208

14.

15.

16.

17.

18.

19.

20.

21.

22.

23.

24.

25.

26.

27.

28.


Chemical Society, Oxford University Press: England, 1998; 321–338. Saiz-Jimenez, C. The chemical structure of humic substances: recent advances. In Humic Substances in Terrestrial Ecosystems; Piccolo, A., Ed.; Elsevier: Amsterdam, The Netherlands, 1996; 1–44. Knicker, H.; Nanny, M.A. Nuclear magnetic resonance spectroscopy. basic theory and background. In NMR Spectroscopy in Environmental Science and Technology; Nanny, M.A., Minear, R.A., Leenheer, J.A., Eds.; Oxford University Press: London, England, 1997; 3–15. Skjemstad, J.O.; Clarke, P.; Golchin, A.; Oades, J.M. Characterisation of soil organic matter by solid-state 13 C nmr spectroscopy. In Driven by Nature: Plant Litter Quality and Decomposition; Giller, K.E., Ed.; CAB International: Wallingford, England, 1997; 253–271. Pfeffer, P.E.; Gerasimowicz, W.V. Nuclear Magnetic Resonance in Agriculture; CRC Press: Boca Raton, FL, 1989. Skjemstad, J.O.; Clarke, P.; Taylor, J.A.; Oades, J.M.; McClure, S.G. The chemistry and nature of protected carbon in soil. Aust. J. Soil Res. 1996, 34, 251–271. Piccolo, A.; Conte, P. Advances in nuclear magnetic resonance and infrared spectroscopies of soil organic particles. In Structure and Surface Reactions of Soil Particles; Huang, P.M., Senesi, N., Buffle, J., Eds.; John Wiley & Sons: Chichester, England, 1998; 183–250. Cheshire, M.V.; Senesi, N. Electron spin resonance spectroscopy of organic and mineral soil particles. In Structure and Surface Reactions of Soil Particles; Huang, P.M., Senesi, N., Buffle, J., Eds.; Wiley: Chichester, England, 1998; 325–376. Huang, Y.; Eglinton, G.; VanderHage, E.R.E.; Boon, J.J.; Bol, R.; Ineson, P. Dissolved organic matter in grass upland soil horizons studied by analytical pyrolysis techniques. Eur. J. Soil Sci. 1998, 49, 1–15. Parsons, J.W. Chemistry and distribution of amino sugars in soils and soil organisms. In Soil Biochemistry; Paul, E.A., Ladd, J.N., Eds.; Marcel Dekker: New York, NY, 1981; Vol. 5, 197–227. Haider, K. Problems related to the humification processes in soils of the temperate climate. In Soil Biochemistry; Bollag, J.-M., Stotzky, G., Eds.; Marcel Dekker: New York, NY, 1992; Vol. 7, 55–94. Shevchenko, S.M.; Bailey, G.W. Life after death: lignin– humic relationships reexamined. Critical Rev. Environ. Sci. Technol. 1996, 26, 95–153. Ko¨gel-Knabner, I.; Ziegler, F.; Riederer, M.; Zech, W. Distribution and decomposition pattern of cutin and suberin in forest soils. Z. Pflanzenerna¨hr. Bodenk. 1989, 152, 409–413. Riederer, M.; Matzke, K.; Ziegler, F.; Ko¨gel-Knanber, I. Inventories and decomposition of the lipid plant biopolymers cutin and suberin in temperate forest soils. Org. Geochem. 1993, 20, 1063–1076. Tegelaar, E.W.; de Leeuw, J.W.; Saiz-Jimenez, C. Possible origin of aliphatic moieties in humic substances. Sci. Total Environ. 1989, 81=82, 1–17. Lichfouse, E.; Chenu, C.; Baudin, F.; Leblond, C.; da Silva, M.; Behar, F.; Derenne, S.; Largeau, C.;


29.

30.

31.

32.

33.

34.

35. 36.

37.

38.

39.

40.

41.

42.

43.

Wehrung, P.; Albrecht, P. A novel pathway of soil organic matter formation by selective preservation of resistant straight-chain biopolymers: chemical and isotopic evidence. Org. Geochem. 1998, 28, 411–415. Waksman, S.A. Humus, Origin, Chemical Composition and Importance in Nature; Ballie´re, Tindall & Cox: London, England, 1938. Aiken, G.R.; McKnight, D.M.; Wershaw, R.L.; MacCarthy, P. Humic Substances in Soil, Sediment and Water; John Wiley & Sons: New York, NY, 1985. Hayes, M.H.B.; MacCarthy, P.; Malcolm, R.; Swift, R.S. Humic Substances II. In Search of Structure; Wiley Interscience: Chichester, England, 1989. Hedges, J.I.; Eglinton, G.; Hatcher, P.G.; Kirchman, D.L.; Arnosti, C.; Derenne, S.; Evershed, R.P.; Ko¨gel-Knabner, I.; de Leeuw, J.W.; Littke, R.; Michaelis, W.; Rullko¨tter, J. The molecularlyuncharacterized component of nonliving organic matter in natural environments. Org. Geochem. 2000, 31, 945–958. Schulten, H.-R.; Leinweber, P.; Reuter, G. Initial formation of soil organic matter from grass residues in a long-term experiment. Biol. Fertil. Soils 1992, 14, 237–245. Stevenson, F.J. Humus Chemistry, Genesis, Composition, Reactions, 2nd Ed.; John Wiley & Sons: New York, NY, 1994. Cheshire, M.V. Nature and Origin of Carbohydrates; Academic Press: London, England, 1979. Chen, C.-N.; Shufeldt, R.C.; Stevenson, F.J. Amino acid analysis of soils and sediments: extraction and desalting. Soil Biol. Biochem. 1975, 7, 143–151. Amelung, W.; Zhang, X. Determination of amino acid enantiomers in soils. Soil Biol. Biochem. 2001, 33, 553–562. Amelung, W. Methods using amino sugars as markers for microbial residues in soil. In Assessment Methods for Soil Carbon; Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; Lewis Publishers: Boca Raton, FL, 2001; 233–272. Dinel, H.; Schnitzer, M.; Mehuys, G.R. Soil lipids origin, nature, content, decomposition and effect on soil physical properties. In Soil Biochemistry; Bollag, J.-M., Stotzky, G., Eds.; Marcel Dekker: New York, NY, 1990; Vol. 5, 397–429. Capriel, P.; Beck, T.; Borchert, H.; Ha¨rter, P. Relationship between soil aliphatic fraction extracted with supercritical hexane, soil microbial biomass, and soil aggregate stability. Soil Sci. Soc. Am. J. 1990, 54, 415–420. Ertel, J.R.; Hedges, J.I. The lignin component of humic substances: distribution among soil and sedimentary humic, fulvic, and base-insoluble fractions. Geochim. Cosmochim. Acta 1984, 48, 2065–2074. Ko¨gel-Knabner, I.; Zech, W.; Hatcher, P.G. Chemical structural studies of forest soil humic acids: aromatic carbon fraction. Soil Sci. Soc. Am. J. 1991, 55, 241–247. Gon˜i, M.A.; Hedges, J.I. Potential applications of cutin-derived cuo reaction products for discriminating vascular plant sources in natural environments. Geochim. Cosmochim. Acta 1990, 54, 23,073–23,081.


44. Glaser, B.; Haumaier, L.; Guggenberger, G.; Zech, W. Black carbon in soils: the use of benzencarboxylic acids as specific markers. Org. Geochem. 1998, 29, 811–819. 45. Glaser, B.; Haumaier, L.; Guggenberger, G.; Zech, W. The terra preta phenomenon: a model for sustainable agriculture in the humid tropics. Naturwissenschaften 2001, 88, 37–41. 46. del Rio, J.C.; McKinney, D.E.; Knicker, H.; Nanny, M.A.; Minard, R.D.; Hatcher, P.G. Structural characterization of bio- and geo-macromolecules by off-line thermochemolysis with tetramethylammonium hydroxide. J. Chromatogr. 1998, 823, 433–448. 47. Filley, T.R.; Minard, R.D.; Hatcher, P.G. Tetramethylammonium hydroxide (TMAH) thermochemolysis:


1209

proposed mechanisms based upon the application of 13 C-labeled TMAH to a synthetic model lignin dimer. Org. Geochem. 1999, 30, 607–621. 48. Preston, C.M. Applications of NMR to soil organic matter analysis: history and prospects. Soil Sci. 1996, 161, 144–166. 49. Preston, C.M. Review of solution NMR of humic substances. In NMR of Humic Substances and Coal; Wershaw, R.L., Mikita, M.A., Eds.; Lewis Publishers: Chelsea, MI, 1987; 3–32. 50. Parfitt, R.L.; Fraser, A.R.; Farmer, V.C. Adsorption on hydrous oxides III. Fulvic acid and humic acid on goethite, gibbsite and imogolite. J. Soil Sci. 1977, 28, 289–296.

Organic Matter Turnover Johan Six Colorado State University, Fort Collins, Colorado, U.S.A.

Julie D. Jastrow Argonne National Laboratory, Argonne, Illinois, U.S.A.

INTRODUCTION Soil organic matter (SOM) is a dynamic entity. The amount (stock) of organic matter in a given soil can increase or decrease depending on numerous factors including climate, vegetation type, nutrient availability, disturbance, land use, and management practices. But even when stocks are at equilibrium, SOM is in a continual state of flux; new inputs cycle—via the process of decomposition—into and through organic matter pools of various qualities and replace materials that are either transferred to other pools or mineralized. For the functioning of a soil ecosystem, this ‘‘turnover’’ of SOM is probably more significant than the sizes of SOM stocks.[1] An understanding of SOM turnover is crucial for quantifying C and nutrient cycles and for determining the quantitative and temporal responses of local, regional, or global C and nutrient budgets to perturbations caused by human activities or climate change.[2]

DEFINITION OF SOIL ORGANIC MATTER TURNOVER The turnover of an element (e.g., C, N, P) in a pool is generally determined by the balance between inputs (I) and outputs (O) of the element to and from the pool Fig. 1. Turnover is most often quantified as the element’s mean residence time (MRT) or its half-life (T1=2). The MRT of an element in a pool is defined as 1) the average time the element resides in the pool at steady state or 2) the average time required to completely renew the content of the pool at steady state. The term half-life is adopted from radioisotope work, where it is defined as the time required for half of a population of elements to disintegrate. Thus, the half-life of SOM is the time required for half of the currently existing stock to decompose. The most common model used to describe the dynamic behavior or turnover of SOM is the firstorder model, which assumes constant zero-order 1210 Copyright © 2006 by Taylor & Francis

input with constant proportional mass loss per unit time[3,4]

@S ¼ I kS; @t

ð1Þ

where S is the SOM stock, t is the time, k is the decomposition rate, and kS is equivalent to output O. Assuming equilibrium (I ¼ O), the MRT can then be calculated as MRT ¼

1 k

ð2Þ

and MRT and T1=2 can be calculated interchangeably with the formula MRT ¼ T1=2 = ln 2

ð3Þ

MEASURING SOIL ORGANIC MATTER TURNOVER Most often the turnover of SOM, more specifically the turnover of SOM-C, is estimated by one of four techniques: 1. Simple first-order modeling 2. 13C natural abundance technique 3. 14C dating technique 4. ‘‘bomb’’ 14C technique. This list does not include tracer studies where a substrate (e.g., plant material) enriched in 13C, 14C, and=or 15 N is added to soil, and its fate is followed over time. Most studies of this type (see Ref.[5] for a review) use the tracers to quantify the short-term (1–5 yr) decomposition rate of freshly added material rather than the long-term turnover of whole-soil C. Eqs. (1) and (2) form the basis for estimates of SOM turnover derived from first-order modeling; the unknown k is calculated as

k ¼

I S

Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001812 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter Turnover

1211

Fig. 1 The turnover of soil organic matter (SOM) is determined by the balance of inputs and outputs. Total SOM consists of many different pools that are turning over at different rates. The mean residence time (MRT) of total SOM is a function of the turnover rates of its constituent pools.

the change (t ¼ 0). An average value of I can then be calculated

by assuming a steady state

@S ¼ 0: @t

I ¼ kSe ;

This approach requires estimates of annual C input rates, which can be assumed to be continuous or discrete.[3] The input can also be written as I ¼ hA where A is the annual addition of C as fresh residue and h (the isohumification coefficient) represents the fraction that, after a rapid initial decomposition of A, remains as the actual annual input to S. An estimate of h is then necessary. A value of 0.3 is commonly used for agricultural crops, but the value can be higher for other materials such as grasses or peat.[6,7] Another approach to estimate k by first-order modeling is ‘‘chronosequence modeling.’’[8] An increase (or decrease) in C across a chronosequence of change in vegetation, land use, or management practice can be fitted to a first-order model Se S0 kt e S ¼ Se 1 Se

MRT ¼

which is equivalent to S ¼ S0 þ ðSe S0 Þð1 e kt Þ

ð4Þ

where t is the time since the change, Se is the C content at equilibrium, and S0 is the initial C content before


but in this case I represents annual inputs of new SOM (hA) rather than inputs of fresh litter or detritus. This approach is also used for chronosequences of primary succession (e.g., on glacial moraines, volcanic deposits, river terraces, dune systems), in this case S0 ¼ 0.[4] The 13C natural abundance technique relies on 1) the difference in 13C natural abundance between plants with different photosynthetic pathways [Calvin cycle (C3 plants) vs. Hatch–Slack cycle (C4 plants)]; and 2) the assumption that the 13C natural abundance signature of SOM is identical to the 13C natural abundance signature of the plants from which it is derived.[9] Thus, where a change in vegetation type has occurred at some known point in time, the rate of loss of the C derived from the original vegetation and the incorporation of C derived from the new vegetation can be inferred from the resulting change in the 13C natural abundance signature of the soil. The turnover of C derived from the original vegetation is then calculated by using the first-order decay model 1 t ¼ k lnðSt =S0 Þ

ð5Þ

where t is the time since conversion, St is the C content derived from original vegetation at time t, and S0 is the C content at t ¼ 0. For further details on the technique see Refs.[9,10].

1212


The presence of 14C with a half-life of 5570 yr in plants and the transformation of this 14C into SOM with little isotopic discrimination allows the SOM to be dated, providing an estimate of the age of the SOM. The 14C dating technique is applicable within a time frame of 200–40,000 yr; samples with an age less than 200 yr are designated as modern (See Ref.[11] for further details of the methodology.) Thermonuclear bomb tests in the 1950s and 1960s caused the atmospheric 14C content to increase sharply and then to fall drastically after the tests were halted. This sequence of events created an in situ tracer experiment; the incorporation of bomb-produced radiocarbon into SOM after the tests stopped allows estimates of the turnover of SOM. Further details of the technique are described in Refs.[2,12,13] RANGE AND VARIATION IN ESTIMATES OF TOTAL SOIL ORGANIC MATTER TURNOVER Comparisons of MRT values estimated by the four methods previously described (see, also, Table 1) reveal a wide range of MRTs. Although variations within each method are attributable to differences in vegetation, climate, soil type, and other factors, the largest variations in observed MRTs are method dependent. For example, MRTs estimated by simple first-order modeling and 13C natural abundance are generally

smaller by an order of magnitude than MRTs estimated by radiocarbon dating, because of the different time scales that the two methods measure. The 13C method is generally used in medium-term observations or experiments (5–50 yr); hence, this method gives an estimate of turnover dominated by relatively recent inputs and C pools that cycle within the time frame of the experiment. In contrast, the oldest and most recalcitrant C pools dominate estimates by radiocarbon dating because of the long-term time frame (200–40,000 yr) that this method measures.[11]

FACTORS CONTROLLING SOIL ORGANIC MATTER TURNOVER Primary production (specifically, the rate of organic matter transfer below-ground) and soil microbial activity (specifically, the rates of SOM transformation and decay) are recognized as the overall biological processes governing inputs and outputs and, hence, SOM turnover. These two processes (and the balance between them) are controlled by complex underlying biotic and abiotic interactions and feedbacks, most of which can be tied in some way to the state factor model of soil formation.[4] Climate (especially temperature and precipitation) constrains both production and decomposition of SOM. Vegetation type affects

Table 1 Range and average mean residence times (MRTs) of total soil organic C in various ecosystem types as estimated by four different methods MRT (yr) Method and ecosystem First-order modeling Cultivated systems and recovering grassland or woodland systems

Sites and sourcesa 7=7

Highb

Average SEc

[14]

102[15]

67 12

Lowb 15

13

C natural abundance Cultivated systems Pasture systems Forest systems

20=10 12=10 2=2

18[16] 17[18] 18[20]

165[17] 102[19] 25[21]

61 9 38 7 22 4

Radiocarbon agingd Cultivated systems Grassland systems Forest systems

21=8e 4=3f 4=3

327[22] Modern[23] 422[22]

1770[23] 1040[24] 1550[25]

880 105 —g 1005 184

‘‘Bomb’’ 14C analysis Cultivated systems Forest and grassland systems

1=1 14=12

1863[13] 36[26]

1863[13] 1542[27]

1863g 535 134

a

First value indicates the number of sites used to calculate average MRT values; second value indicates the number of literature sources surveyed (i.e., some sources provided data for multiple sites). b Number in parentheses indicates reference to literature. c SE, standard error. d Values presented in MRT columns for this technique are radiocarbon ages in years B.P. e Includes two sites dating as ‘‘modern.’’ f Includes three sites dating as ‘‘modern.’’ g Only one value available.



1213

production rates and the types and quality of organic inputs (e.g., below- vs. above-ground, amounts of structural tissue, C=N and lignin=N ratios), as well as the rates of water and nutrient uptake—all of which, in turn, influence decomposition rates. The types, populations, and activities of soil biota control decomposition and nutrient cycling=availability and hence influence vegetative productivity. Parent material affects SOM turnover as soil type, mineralogy, texture, and structure influence pH, water and nutrient supply, aeration, and the habitat for soil biota, among other factors. Topography modifies climate, vegetation type, and soil type on the landscape scale and exerts finerscale effects on temperature, soil moisture, and texture. Lastly, time affects whether inputs and outputs are at equilibrium, and temporal scale influences the relative importance of various state factor effects on production and decomposition. Disturbance or management practices also exert considerable influence on SOM turnover via direct effects on inputs and outputs and through indirect effects on the factors controlling these fluxes. An example of management effects on MRT is illustrated in Table 2; in most cases, the MRT of whole-soil C is significantly longer under no tillage agriculture than under conventional tillage practices.

TURNOVER OF DIFFERENT SOIL ORGANIC MATTER POOLS The previous discussion is focused on the turnover and MRT of whole-soil C; hence, it treats SOM as a single, homogeneous reservoir. But, in fact, SOM is a

heterogeneous mixture consisting of plant, animal, and microbial materials in all stages of decay combined with a variety of decomposition products of different ages and levels of complexity. Thus, the turnover of these components varies continuously, and any estimate of MRT for SOM as a whole merely represents an overall average value (Fig. 1). Although average MRTs are useful for general comparisons of sites or the effects of different management practices, they can be misleading because soils with similar average MRTs can have very different distributions of organic matter among pools with fast, slow, and intermediate turnover rates.[2,31] Simulation models that account for variations in turnover rates for different SOM pools are now used to generate more realistic descriptions of SOM dynamics. A few models represent decomposition as a continuum, with each input cohort following a pattern of increasing resistance to decay,[32] but most models are multicompartmental, with several organic matter pools (often 3–5) that are kinetically defined with differing turnover rates. For example, the CENTURY SOM model[33] divides soil C into active, slow, and passive pools, with MRTs of 1.5, 25, and about 1000 yr, respectively, and separates plant inputs into metabolic (readily decomposable; MRT of 0.1–1 yr) and structural (difficult to decompose; MRT of 1–5 yr) pools as a function of lignin : N ratio. Even though compartmental models are reasonably good at simulating changes in SOM, the compartments are conceptual in nature, and thus it has been difficult to relate them to functionally meaningful pools or experimentally verifiable fractions.[34,35] The use of isotopic techniques to analytically determine the MRTs of physically and chemically

Table 2 Effect of tillage practices on mean residence time (MRT) of total soil organic C estimated by the 13C natural abundance technique Cropping systema

Site (Ref.) [28]

Wheat–fallow (NT)

Sidney, NE

Depth (cm)

tb (yr)

MRT (yr)

0–20

26

73

0–20

5

26

0–30

17

127

0–30

11

118

Wheat–fallow (CT) Delhi, Ont.[29]

Corn (NT)

44

Corn (CT) Boigneville, France[16]

Corn (NT)

14

Corn (CT) Rosemount, MN[30]

Corn (NT, 200 kg N ha1 yr1)

55

Corn (CT, 200 kg N ha1 yr1)

73

Corn (NT, 0 kg N ha1 yr1)

54

Corn (CT, 0 kg N ha1 yr1) Average SE

c

a

NT, no tillage; CT, conventional (moldboard plow) tillage. b Time period of experiment. c SE, standard error.


72 NT

80 19

CT

52 11

1214


Table 3 Mean residence time (MRT) of macro- and microaggregate-associated C estimated by the 13C natural abundance technique

macroaggregates vs. microaggregates show consistently slower turnovers in microaggregates (Table 3). Thus, a much higher proportion of the SOM occluded in microaggregates consists of stabilized materials with relatively long MRTs.

Aggregate size classa

lm

MRT (yr)

Tropical pasture[44]

M m

>200 250 50–250

14 61

Corn[47]

M m

>250 50–250

42 691

Wheat–fallow, no tillage[48]

M m

250–2000 53–250

27 137

Wheat–fallow, conventional tillage[48]

M m

250–2000 53–250

8 79

Average SEb

M m

1. Paul, E.A. Dynamics of organic matter in soils. Plant Soil 1984, 76, 275–285. 2. Trumbore, S.E. Comparison of carbon dynamics in tropical and temperate soils using radiocarbon measurements. Global Biogeochem. Cycles 1993, 7, 275–290. 3. Olson, J.S. Energy storage and the balance of producers and decomposers in ecological systems. Ecology 1963, 44, 322–331. 4. Jenny, H. The Soil Resource—Origin and Behavior; Springer: New York, 1980; 377 pp. 5. Schimel, D.S. Theory and Application of Tracers; Academic Press: San Diego, CA, 1993; 119 pp. 6. Buyanovsky, G.A.; Kucera, C.L.; Wagner, G.H. Comparative analyses of carbon dynamics in native and cultivated ecosystems. Ecology 1987, 68, 2023–2031. 7. Jenkinson, D.S. The turnover of organic carbon and nitrogen in soil. Phil. Trans. R. Soc. Lond. Ser. B 1990, 329, 361–368. 8. Jastrow, J.D. Soil aggregate formation and the accrual of particulate and mineral-associated organic matter. Soil Biol. Biochem. 1996, 28, 665–676. 9. Cerri, C.; Feller, C.; Balesdent, J.; Victoria, R.; Plenecassagne, A. Application du tracage isotopique natural en 13 C a l’etude de la dynamique de la matiere oganique dans les sols. C.R. Acad. Sci. Paris Ser. II 1985, 300, 423–428. 10. Balesdent, J.; Mariotti, A. Measurement of soil organic matter turnover using 13C natural abundance. In Mass Spectrometry of Soils; Boutton, T.W., Yamasaki, S., Eds.; Marcel Dekker: New York, 1996; 83–111. 11. Goh, K.M. Carbon dating. In Carbon Isotope Techniques; Coleman, D.C., Fry, B., Eds.; Academic Press: San Diego, CA, 1991; 125–145. 12. Goh, K.M. Bomb carbon. In Carbon Isotope Techniques; Coleman, D.C., Fry, B., Eds.; Academic Press: San Diego, CA, 1991; 147–151. 13. Harrison, K.G.; Broecker, W.S.; Bonani, G. The effect of changing land use on soil radiocarbon. Science 1993, 262, 725–726. 14. Hendrix, P.F. Long-term patterns of plant production and soil carbon dynamics in a georgia piedmont agroecosystem. In Soil Organic Matter in Temperate Agroecosystems: Long-Term Experiments in North America; Paul, E.A., Paustian, K., Elliott, E.T., Cole, C.V., Eds.; CRC Press: Boca Raton, FL, 1997; 235–245. 15. Buyanovsky, G.A.; Brown, J.R.; Wagner, G.H. Sanborn field: effect of 100 years of cropping on soil parameters influencing productivity. In Soil Organic Matter in Temperate Agroecosystems: Long-Term Experiments in North America; Paul, E.A., Paustian, K., Elliott, E.T., Cole, C.V., Eds.; CRC Press: Boca Raton, FL, 1997; 205–225.

Ecosystem (Ref.)

a

42 18 209 95

M, macroaggregate; m, microaggregate. SE, standard error.

b

separated SOM fractions has demonstrated the existence of various turnover rates for different pools. For example, low-density SOM (except for charcoal) invariably turns over faster than high-density, mineral-associated SOM, and hydrolyzable SOM turns over faster than nonhydrolyzable residues.[36,37] The MRTs of primary organomineral associations generally increase with decreasing particle size, although there are exceptions (particularly among fine gradations of silt- and clay-sized particles) that have been variously related to climate, clay mineralogy, and fractionation methodology.[34,38,39] For a given set of biotic and abiotic conditions, the turnover of different SOM pools depends mechanistically on the quality and biochemical recalcitrance of the organic matter and its accessibility to decomposers. With other factors equal, clay soils retain more SOM with longer MRTs than do sandy soils.[40] Readily decomposable materials can become chemically protected from decomposition by association with clay minerals and by sorption to humic colloids.[38,41] Clay mineralogy also plays an important role. For example, montmorillonitic clays and allophanes generally afford more protection than illites and kaolinites.[42] In addition, the spatial location of SOM within the soil matrix determines its physical accessibility to decomposers. Relatively labile material may become physically protected by incorporation into soil aggregates[43] or by deposition in micropores inaccessible even to bacteria. Studies of the average MRTs of organic matter in



16. Balesdent, J.; Mariotti, A.; Boisgontier, D. Effect of tillage on soil organic carbon mineralization estimated from 13C abundance in maize fields. J. Soil Sci. 1990, 41, 587–596. 17. Vitorello, V.A.; Cerri, C.C.; Andreux, F.; Feller, C.; Victoria, R.L. Organic matter and natural carbon-13 distributions in forested and cultivated oxisols. Soil Sci. Soc. Am. J. 1989, 53, 773–778. 18. Desjardins, T.; Andreux, F.; Volkoff, B.; Cerri, C.C. Organic carbon and 13C contents in soils and soil sizefractions, and their changes due to deforestation and pasture installation in Eastern Amazonia. Geoderma 1994, 61, 103–118. 19. Jastrow, J.D.; Boutton, T.W.; Miller, R.M. Carbon dynamics of aggregate-associated organic matter estimated by carbon-13 natural abundance. Soil Sci. Soc. Am. J. 1996, 60, 801–807. 20. Martin, A.; Mariotti, A.; Balesdent, J.; Lavelle, P.; Vuattoux, R. Estimate of organic matter turnover rate in a savanna soil by 13C natural abundance measurements. Soil Biol. Biochem. 1990, 22, 517–523. 21. Trouve, C.; Mariotti, A.; Schwartz, D.; Guillet, B. Soil organic carbon dynamics under eucalyptus and pinus planted on savannas in the congo. Soil Biol. Biochem. 1994, 26, 287–295. 22. Paul, E.A.; Collins, H.P.; Leavitt, S.W. Dynamics of resistant soil carbon of midwestern agricultural soils measured by naturally-occurring 14C abundance. Geoderma 2001, 104, 239–256. 23. Paul, E.A.; Follett, R.F.; Leavitt, S.W.; Halvorson, A.; Peterson, G.A.; Lyon, D.J. Radiocarbon dating for determination of soil organic matter pool sizes and dynamics. Soil Sci. Soc. Am. J. 1997, 61, 1058–1067. 24. Jenkinson, D.S.; Harkness, D.D.; Vance, E.D.; Adams, D.E.; Harrison, A.F. Calculating net primary production and annual input of organic matter to soil from the amount and radiocarbon content of soil organic matter. Soil Biol. Biochem. 1992, 24, 295–308. 25. Trumbore, S.E.; Bonani, G.; Wolfli, W. The rates of carbon cycling in several soils from AMS 14C measurement of fractionated soil organic matter. In Soils and the Greenhouse Effect; Bouwman, A.F., Ed.; Wiley: London, 1990; 407–414. 26. O’Brien, B.J. Soil organic carbon fluxes and turnover rates estimated from radiocarbon enrichments. Soil Biol. Biochem. 1984, 16, 115–120. 27. Bol, R.A.; Harkness, D.D.; Huang, Y.; Howard, D.M. The influence of soil processes on carbon isotope distribution and turnover in the british uplands. Eur. J. Soil Sci. 1999, 50, 41–51. 28. Six, J.; Elliott, E.T.; Paustian, K.; Doran, J.W. Aggregation and soil organic matter accumulation in cultivated and native grassland soils. Soil Sci. Soc. Am. J. 1998, 62, 1367–1377. 29. Ryan, M.C.; Aravena, R.; Gillham, R.W. The use of 13C natural abundance to investigate the turnover of the microbial biomass and active fractions of soil organic matter under two tillage treatments. In Soils and Global Change; Lal, R., Kimble, J., Levine, E., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1995; 351–360. 30. Clapp, C.E.; Allmaras, R.R.; Layese, M.F.; Linden, D.R.; Dowdy, R.H. Soil organic carbon and 13C


1215

31.

32.

33.

34.

35.

36.

37.

38.

39. 40.

41.

42.

43. 44.

45.

46.

47.

48.

abundance as related to tillage, crop residue, and nitrogen fertilization under continuous corn management in Minnesota. Soil Till. Res. 2000, 55, 127–142. Davidson, E.A.; Trumbore, S.E.; Amundson, R. Soil warming and organic carbon content. Nature 2000, 408, 789–790. ˚ gren, G.I.; Bosatta, E. Theoretical analysis of the longA term dynamics of carbon and nitrogen in soils. Ecology 1987, 68, 1181–1189. Parton, W.J.; Schimel, D.S.; Cole, C.V.; Ojima, D.S. Analysis of factors controlling soil organic matter levels in great plains grasslands. Soil Sci. Soc. Am. J. 1987, 51, 1173–1179. Balesdent, J. The significance of organic separates to carbon dynamics and its modeling in some cultivated soils. Eur. J. Soil Sci. 1996, 47, 485–493. Christensen, B.T. Matching measurable soil organic matter fractions with conceptual pools in simulation models of carbon turnover: revision of model structure. In Evaluation of Soil Organic Matter Models; Powlson, D.S., Smith, P., Smith, J.U., Eds.; Springer: Berlin, 1996; 143–159. Martel, Y.A.; Paul, E.A. The use of radiocarbon dating of organic matter in the study of soil genesis. Soil Sci. Soc. Am. Proc. 1974, 38, 501–506. Trumbore, S.E.; Chadwick, O.A.; Amundson, R. Rapid exchange between soil carbon and atmospheric carbon dioxide driven by temperature change. Science 1996, 272, 393–396. Christensen, B.T. Physical fractionation of soil and organic matter in primary particle size and density separates. Adv. Soil Sci. 1992, 20, 1–90. Feller, C.; Beare, M.H. Physical control of soil organic matter dynamics in the tropics. Geoderma 1997, 79, 69–116. Sorensen, L.H. The influence of clay on the rate of decay of amino acid metabolites synthesized in soils during decomposition of cellulose. Soil. Biol. Biochem. 1974, 7, 171–177. Jenkinson, D.S. Soil organic matter and its dynamics. In Russell’s Soil Conditions and Plant Growth; Wild, A., Ed.; Wiley: New York, 1988; 564–607. Dalal, R.C.; Bridge, B.J. Aggregation and organic matter storage in sub-humid and semi-arid soils. In Structure and Organic Matter Storage in Agricultural Soils; Carter, M.R., Stewart, B.A., Eds.; CRC Press: Boca Raton, FL, 1996; 263–307. Tisdall, J.M.; Oades, J.M. Organic matter and waterstable aggregates in soils. J. Soil Sci. 1982, 33, 141–163. Skjemstad, J.O.; Le Feuvre, R.P.; Prebble, R.E. Turnover of soil organic matter under pasture as determined by 13C natural abundance. Aust. J. Soil Res. 1990, 28, 267–276. Buyanovsky, G.A.; Aslam, M.; Wagner, G.H. Carbon turnover in soil physical fractions. Soil Sci. Soc. Am. J. 1994, 58, 1167–1173. Monreal, C.M.; Schulten, H.R.; Kodama, H. Age, turnover and molecular diversity of soil organic matter in aggregates of a gleysol. Can. J. Soil Sci. 1997, 77, 379–388. Angers, D.A.; Giroux, M. Recently deposited organic matter in soil water-stable aggregates. Soil Sci. Soc. Am. J. 1996, 60, 1547–1551. Six, J.; Elliott, E.T.; Paustian, K. Aggregate and soil organic matter dynamics under conventional and no-tillage systems. Soil Sci. Soc. Am. J. 1999, 63, 1350–1358.

Organic Matter: Global Distribution in World Ecosystems Wilfred M. Post Oak Ridge National Laboratory, Oak Ridge, Tennessee, U.S.A.

INTRODUCTION Globally, the amount of organic matter in soils, commonly represented by the mass of carbon, is estimated to be 1200–1500 Pg C (1 Pg C ¼ 1015 g carbon) in the top 1 m of soil.[1,2] This is 2–3 times larger than the amount of organic matter in living organisms in all terrestrial ecosystems.[1] The exact ratio between living and dead organic matter in terrestrial ecosystems varies, depending on the ecosystem. The amount of carbon stored in soil is determined by the balance of two biotic processes—the productivity of terrestrial vegetation and the decomposition of organic matter. Each of these processes has strong physical and biological controlling factors. These include climate; soil chemical, physical, and biological properties; and vegetation composition. Interactions among these controlling factors are of particular importance. These biological and physical factors are the same as the ones that influence the above ground structure and composition of terrestrial ecosystems, so there are strong correspondences between soil organic matter content and ecosystem type.

ORGANIC MATTER INPUTS Quantity The amount of carbon stored in soils is to a great extent determined by the rate of organic matter input through litterfall, root exudates, and root turnover. The main factors that influence vegetation production are suitable temperatures for photosynthesis, available soil moisture for evapotranspiration, and rates of CO2 and H2O exchange. Dry and=or cold climates support low vegetation production rates and soils under such climates have low organic matter contents. Where climates are warm and moist, vegetation production is high and soil organic matter contents are correspondingly high. Fig. 1 shows the striking correspondence between soil organic matter content and general climate measurements that results from the relationship between vegetation production and suitable moisture and temperature conditions. Vegetation production depends not only on climate but also on nutrient supply from decomposition and 1216 Copyright © 2006 by Taylor & Francis

geochemical weathering. Walker and Adams[6] hypothesized that the level of available phosphorus during the course of soil development is the primary determinant of terrestrial net primary production. Numerous workers have examined this hypothesis. Tiessen, Stewart, and Cole[7] and Roberts, Stewart, and Bettany[8] found that available phosphorus explained about one-fourth of the variance in soil organic matter in many different soil orders. The relationship between phosphorus and carbon is strongest during the aggrading stage of vegetation–soil system development.[9] Initially, the production of acidic products by pioneer vegetation promotes the release of phosphorus by weathering of parent material. Organic matter builds up in the soil, increasing the storage of phosphorus in decomposing organic compounds. Nitrogen fixing bacteria populations, which depend on a supply of organic carbon and available phosphorus, can grow to meet ecosystem demands for nitrogen. Plant growth is enhanced by this increasing nitrogen and phosphorus cycling, resulting in increased rates of weathering. This process continues until the vegetation is constrained by other factors affecting phosphorus availability: Leaching losses become larger than the weathering inputs;[10] or an increasing fraction of the phosphorus becomes unavailable by adsorption or precipitation with secondary minerals;[11] or nitrogen availability (denitrification or leaching is affected) reaching or exceeding nitrogen inputs and fixation.[12] In mature soils, net primary production is more likely to be limited by nitrogen. Availability of other nutrients that are largely derived from parent materials, such as most base cations, may also influence soil organic matter accumulation during early soil development.[13] Soils derived from base cation rich volcanic parent materials (Andisols) have much higher carbon contents on average than soils from other parent materials.[4] Species Composition Biotic factors, in particular plant species composition, also affect soil organic matter dynamics. Production and decomposition rates are to some degree controlled by species composition. Each terrestrial plant species produces different amounts and chemical compositions of leaves, roots, branches, and wood of varying decomposability. This range of decomposability may be Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001792 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organic Matter: Global Distribution in World Ecosystems

1217

Fig. 1 Contours of soil carbon density (kg m2) plotted on Holdridge diagram[3] for world life-zone classification. Values of biotemperature and precipitation uniquely determine a life zone and associated vegetation. Contour lines for mean soil carbon content in the surface meter of soil are determined from data derived from over 3000 soil profiles.[4,5]

summarized by the lignin and nitrogen content of the organic material.[14,15] Litter decay rate is inversely related to C : N and lignin:N ratios and positively related to N content. Species with tissues that have low nutrient or high lignin content produce litter that is slow to decay. Nitrogen is made available to plants during the decomposition process. Nitrogen is a limiting element for productivity in most terrestrial ecosystems so the rate at which it is released during decomposition is an important factor in ecosystem production. Thus, the interactions between processes regulating plant populations and their productivity and microbial processes regulating nitrogen availability result in some of the observed variation in soil carbon and nitrogen storage.[16–19]

Placement The deeper that fresh detritus is placed in the soil, the slower it decomposes. This is a result of declining


decomposer activity and increased protection from oxidation with depth in the soil. Prairies have a somewhat lower productivity than forests and produce no slowly decomposing woody material. Nevertheless, prairies have a very high soil organic matter content because prairie grasses allocate twice as much production to belowground roots and tillers than to aboveground leaves.[20] The result is high soil organic matter contents with a uniform distribution in the upper 1 m of soil (Fig. 2). In contrast, a spruce–fir forest contains 50 percent of its soil organic matter in the top 10 cm. There are interesting exceptions to the rule that above-=belowground plant allocation determines soil organic matter distribution patterns in soil. Tropical moist forest soils show a uniform depth distribution similar to the depth distribution of temperate grasslands, however, in tropical forests this is largely due to a long-term accumulation of recalcitrant organic materials at lower depths in the soil rather than increased allocation to roots. Alpine tundra soils support a largely herbaceous flora but show a similar

1218


and as precipitation increases in the warm temperate, subtropical, and tropical life zones. The combined influence of temperature and precipitation is presented by the third axis of the Holdridge diagram (Fig. 1) as the ratio of potential evapotranspiration (PET) to annual precipitation. When this ratio is less than 1.0, rainfall exceeds PET and vice versa. Life zones bordering the line with the PET is equal to precipitation (PET ratio ¼ 1.0) have soil carbon contents around 10 kg m2 except in warm temperate and subtropical zones where strong seasonality limits production, but decomposition conditions are favorable for most of the year. Soil carbon content increases as the PET ratio decreases indicating that productivity increases faster than the rate of decomposition with increasing moisture availability.

Organic Matter Quality

Fig. 2 Cumulative carbon storage as a function of depth for four ecosystems. Refer text for explanation of these patterns. (From Ref.[4].)

depth distribution as forest soils because of inhibition of surface litter decomposition by low temperatures and high water saturation.

DECOMPOSITION Climate Organic matter decay rates can be related to environmental parameters such as temperature and soil moisture. Climatic indices that correlate well with decay rates include plant moisture and temperature indices,[21,22] linear combinations of temperature and rainfall,[23] and actual evapotranspiration.[15] Warm temperatures and available soil moisture enhance microbial, and micro- and macro-invertebrate activity. These environmental conditions are also correlated with plant production. As a result, the amount of organic matter present in soil is highest in vegetation types with the highest rates of organic matter production. These are ones found in the warm, moist climate regions. The contours of soil carbon density displayed in Fig. 1 reflect the balance of input by vegetation production and loss from decomposition imposed by climate. Soil carbon content increases from lower left to the upper right in Fig. 1 as the temperature decreases in the cool temperate, boreal, and sub-polar life zones


On global scale, climate may be the most important factor controlling decay rates, but within a given region, substrate chemistry is the more important factor.[15,24,25] Decay rate is often negatively related to substrate C : N ratio. Litter C : N is initially much greater than microbial C : N but approaches microbial C : N as the microbes release the carbon as CO2 while taking up nitrogen (nitrogen immobilization). The further the initial litter C : N is from microbial C : N, the slower the decay rate. Lignin content or lignin: N ratios may be better predictors of decay rates because lignin itself is difficult to decompose, and it shields nitrogen and other more easily degraded chemical fractions from microbes. Concise and simple models of decay rate are based on a combination of chemical and climatic indices. The effect of litter quality on soil organic matter content is most dramatically expressed in Podzols (Spodosol in the United States Department of Agriculture classification). These occur over large areas in boreal zones dominated by evergreen conifers, but often occur in other regions on shallow or sandy soils. Low nitrogen content of organic matter inputs and cool temperatures reduce decomposition and soil animal activity. As a result, large surface organic matter accumulations occur over a thin A horizon. Low temperatures combined with leaching of organic acids result in podsolization as the predominant soil-forming process. Leaching of iron, aluminum oxides, and organic matter result in a distinct E horizon near the surface where these materials are removed and deposited in the B horizon. If the surface organic layers are included, these soils can have substantial organic matter contents, exceeding the expected amount for the climate conditions. Batjes[2] gives an average value for Podzols of 24.2 kg m2 for the surface meter which is


Table 1 Mean organic carbon contents (kg m2) by FAO–UNESCO soil units to 1 m depth Soil unit

Mean C (kg m2)

Acrisols

9.4

Cambisols

9.6

Chernozems Podzoluvisols

12.5 7.3

Ferrasols

10.7

Gleysols

13.1

Phaeozems

14.6

Fluvisols

9.3

Kastanozems

9.6

Luvisols

6.5

Greyzems

19.7

Nitosols

8.4

Histosols

77.6

Podzols

24.2

Arenosols

3.1

Regosols

5.0

Solonetz

6.2

Andisols

25.4

Vertisols

11.1

Planosols

7.7

Xerosols

4.8

Yermosols

3.0

Solochaks

4.2

These soil units generally span a wide range of climate conditions and therefore present a different view of soil organic matter content based on additional soil factors. In particular, the high C content of Podzols, Histosols, and Andisols is apparent. Refer text for additional explanation of biological, chemical and physical factors responsible. (From Ref.[2].)

considerably above the mean for most other soil types (see Table 1). SIGNIFICANT PHYSICAL AND CHEMICAL INFLUENCES There are several notable exceptions to the climatebased explanation of variation in soil carbon content. There are two in particular that have lower rates of decomposition and therefore higher accumulations of organic matter than expected (Table 1). These include Histosols due to hydrological conditions and Andisols due to parent material chemical effects. Histosols In landscape positions where water accumulates at or above the surface of the soil for an appreciable part


1219

of the growing season, decomposition can be reduced to such an extent that large amounts of undecomposed organic matter can accumulate. This soil type is called a Histosol and can be found in any region in wetlands where decomposition is restricted. The soil-surface of mature or old-growth boreal forests over shallow water tables are often covered with Sphagnum moss which may also lead to development of Histosols. Histosols with the largest areas and thickest accumulations occur in lowland tundra where a mixture of sedges, lichens, and mosses grow at the northern limit of vegetation in the northern hemisphere. Production, decomposition, and evaporation are limited by low temperatures and water-saturated soils. In these cold regions, deeper layers may freeze and and not become thawed during the short growing season (permafrost). As a result, Histosols have carbon contents over 70 kg m2 in the surface meter (Table 1). Some regions have been accumulating organic matter since the last glacial period without any substantial decomposition. Histosols in such regions may be several meters thick and contain over 250 kg C m2.[2] Globally it is estimated that boreal and sub-arctic Histosols contain 455 PgC that has accumulated during the postglacial period.[26] Andisols Andisols form on young volcanic stone (basalt lava) rich in nutrients and alkaline. Andisols are weakly weathered soils associated with pyroclastic parent materials that are rich in allophane, ferrihydrite, and other minerals that readily form complexes with humus molecules. These chemical constituents provide conditions promoting high vegetation production and also the retention of organic matter in soil. As a result, Andisols typically have higher soil carbon contents (25.4 kg m2, Table 1) than soils with the same environmental conditions but different parent materials.

CONCLUSIONS Over long periods of time, organic matter in soils is the result of climatic, biological, and geological factors. These factors are not independent. In particular there exists a strong relationship between climate and vegetation type. In Fig. 1, the Holdridge climate based life zones have names that depict the dominant vegetation of climates. Jobba´gy and Jackson[27] provide a summary of soil data based on biomes that demonstrates similar soil carbon distribution as that based on climate (Table 2). Over shorter periods of time soil carbon varies with vegetation disturbances and changes in land use

1220


Table 2 Mean organic carbon content (kg m2) by biome to 1 m depth Biome Boreal forest

Mean C (kg m2) 9.3

Crops

11.2

Deserts

6.2

Sclerophyllous shrubs

8.9

Temperate deciduous forest

17.4

Temperate evergreen forest

14.5

Temperate grassland

11.7

Tropical deciduous forest

15.8

Tropical evergreen forest

18.6

Tropical grassland=savanna

13.2

Tundra

14.2

Biome classification is based on Whittaker.[28] (From Ref.[27].)

patterns that affect rates of organic matter input and its decomposition. Various land uses result in very rapid declines in soil organic matter from the native condition.[29–32] Losses of 50% in the top 20 cm and 30% for the surface 100 cm are average. Much of this loss in soil organic carbon can be attributed to erosion, reduced inputs of organic matter, increased decomposability of crop residues, and tillage effects that decrease the amount of physical protection to decomposition. Evidence from long-term experiments suggest that C losses due to oxidation and erosion can be reversed with soil management practices that minimize soil disturbance and optimize plant yield through fertilization. These experimental results are believed to apply to large regions and that organic matter is being restored as a result of establishment of perennial vegetation, increased adoption of conservation tillage methods, efficient use of fertilizers, and increased use of high yielding crop varieties.[33,34] Additionally, when agricultural land is no longer used for cultivation and allowed to revert to natural vegetation or replanted to perennial vegetation, soil organic carbon can accumulate by processes that essentially reversing some of the effects responsible for soil organic carbon losses initially—from when the land was converted from perennial vegetation—and return them to typical amounts for the climate, vegetation, landscape position, and parent material conditions.[35,36]

ACKNOWLEDGMENTS Work sponsored by U.S. Department of Energy, Carbon Dioxide Research Program, Environmental Sciences Division, Office of Biological and Environmental Research and performed at Oak Ridge


National Laboratory (ORNL). ORNL is managed by UT-Battelle, LLC, for the U.S. Department of Energy under contract DE-AC05-00OR22725.

REFERENCES 1. Post, W.M.; Peng, T.-H.; Emanuel, W.R.; King, A.W.; Dale, V.H.; DeAngelis, D.L. The global carbon cycle. American Scientist 1990, 78, 310–326. 2. Batjes, N.H. Total carbon and nitrogen in the soils of the world. European Journal of Soil Science 1996, 47, 151–163. 3. Holdridge, L.R. Determination of world plant formations from simple climatic data. Science 1947, 105, 367–368. 4. Zinke, P.J.; Stangenberger, A.G.; Post, W.M.; Emanuel, W.R.; Olson, J.S. Worldwide Organic Soil Carbon and Nitrogen Data; ORNL=TM-8857; Oak Ridge National Laboratory: Oak Ridge, TN, 1984. 5. Post, W.M.; Pastor, J.; Zinke, P.J.; Stangenberger, A.G. Global patterns of soil nitrogen storage. Nature 1985, 317, 613–616. 6. Walker, T.W.; Adams, A.F.R. Studies on soil organic matter: I. Influence of phosphorus content of parent materials on accumulations of carbon, nitrogen, sulfur, and organic phosphorus in grassland soils. Soil Science 1958, 85, 307–318. 7. Tiessen, H.J.; Stewart, W.B.; Cole, C.V. Pathways of phosphorus transformations in soils of differing pedogenesis. Soil Science Society of America Journal 1984, 48, 853–858. 8. Roberts, T.L.; Stewart, J.W.B.; Bettany, J.R. The influence of topography on the distribution of organic and inorganic soil phosphorus across a narrow environmental gradient. Canadian Journal of Soil Science 1985, 65, 651–665. 9. Anderson, D.W. The effect of parent material and soil development on nutrient cycling in temperate ecosystems. Biogeochemistry 1988, 5, 71–97. 10. Jenny, H. The Soil Resource; Springer: Berlin, 1980. 11. Walker, T.W.; Syers, J.K. The fate of phosphorus during pedogenesis. Geoderma 1976, 15, 1–19. 12. Schlesinger, W.H. Biogeochemistry: An Analysis of Global Change; Academic: New York, 1991. 13. Torn, M.S.; Trumbore, S.E.; Chadwick, O.A.; Vitousek, P.M.; Hendricks, D.M. Mineral control of soil organic carbon storage and turnover. Nature 1997, 389, 170–173. 14. Aber, J.D.; Melillo, J.M. Nitrogen immobilization in decaying hardwood leaf litter as a function of initial nitrogen and lignin content. Canadian Journal of Botany 1982, 58, 416–421. 15. Meentemeyer, V. Macroclimate and lignin control oflitter decomposition rates. Ecology 1978, 59, 465–472. 16. Zinke, P.J. The pattern of influence of individual trees on soil properties. Ecology 1962, 42, 130–133. 17. Wedin, D.A.; Tilman, D. Species effects on nitrogen cycling: a test with perennial grasses. Oecologia 1990, 84, 433–441.


18. Hobbie, S.E. Effects of plant species on nutrient cycling. Trends in Ecology and Evolution 1992, 7, 336–339. 19. Hobbie, S.E. Temperature and plant species control over litter decomposition in Alaskan tundra. Ecological Monographs 1996, 66, 503–522. 20. Sims, P.L.; Coupland, R.T. Grassland Ecosystems of the World: Analysis of Grasslands and Their Uses; Coupland, R.T., Ed.; Cambridge University Press: Cambridge, 1979. 21. Olson, J.S. Energy storage and the balance of producers and decomposers in ecological systems. Ecology 1963, 44, 322–331. 22. Fogel, R.; Cromack, K. Effect of habitat and substrate quality on douglas fir litter decomposition in western Oregon. Canadian Journal of Botany 1977, 55, 1632–1640. 23. Pandey, V.; Singh, J.S. Leaf-litter decomposition in an oak–conifer forest in himalaya: the effects of climate and chemical composition. Forestry 1982, 55, 47–59. 24. Flanagan, P.W.; VanCleve, K. Nutrient cycling in relation to decomposition and organic matter quality in tiaga ecosystems. Canadian Journal of Forest Research 1983, 13, 795–817. 25. McClaugherty, C.A.; Pastor, J.; Aber, J.D.; Melillo, J.M. Forest litter decomposition in relation to soil nitrogen dynamics and litter quality. Ecology 1984, 66, 266–275. 26. Gorham, E. Northern peatlands: role in the carbon cycle and probable responses to climatic warming. Ecological Applications 1991, 1, 182–195.


1221

27. Jobba´gy, E.G.; Jackson, R.B. The vertical distribution of organic carbon and its relation to climate and vegetation. Ecological Applications 2000, 10 (2), 423–436. 28. Whittaker, R.H. Communities and Ecosystems; MacMillan: London, 1975. 29. Jenny, H. Factors of Soil Formation; McGraw-Hill: New York, 1941. 30. Davidson, E.A.; Ackerman, I.L. Changes in soil carbon inventories following cultivation of previously untilled soils. Biogeochemistry 1993, 20, 161–193. 31. Mann, L.K. Changes in soil carbon after cultivation. Soil Science 1986, 142, 279–288. 32. Schlesinger, W.H. Changes in soil carbon storage and associated properties with disturbance and recovery. In The Changing Carbon Cycle: A Global Analysis; Trabalka, J.R., Reichle, D.E., Eds.; Springer: New York, 1985. 33. Buyanovsky, G.A.; Wagner, G.H. Carbon cycling in cultivated land and its global significance. Global Change Biology 1998, 4, 131–142. 34. Lal, R.; Kimble, J.M.; Follett, R.F.; Cole, C.V. The Potential of U.S. Cropland to Sequester Carbon and Mitigate the Greenhouse Effect; Ann Arbor Press: Ann Arbor, 1998. 35. Post, W.M.; Kwon, K.C. Soil carbon sequestration and land-use change: processes and potential. Global Change Biology 2000, 6, 317–328. 36. Silver, W.L.; Ostertag, R.; Lugo, A.E. The potential for carbon sequestration through reforestation of abandoned tropical agricultural and pasture lands. Restoration Ecology 2000, 8 (4), 394–407.

Organisms and Soil Food Webs David C. Coleman University of Georgia, Athens, Georgia, U.S.A.

INTRODUCTION Soils may be viewed as the organizing centers for terrestrial ecosystems. This is largely the result of organismal activities in the soil. Major functions such as ecosystem production, respiration and nutrient recycling are controlled by the rates at which nutrients are released by decomposition in the soil and litter horizons. The array of biota, including microbes, microbe-feeding fauna, vegetation, and consumers are all influenced by soil processes, and the organisms in turn have an impact on the soil system. Soils provide a wide range and variety of microhabitats, thus accommodating a very diverse biota. The enormous surface area (hundreds of m2=g of soil) of soil particles, ranges in size classes from clays (0.1– 2 mm in dia), to silts (2–50 mm in dia), and sands (0.05–2 mm in dia). Numerous microbes and microand mesofauna (protozoa and nematodes) exist in water films on these particles, and in or on the surfaces of microaggregates formed from the primary particles.[1] In turn, the more mobile fauna, from collembola and mites (larger mesofauna) to the macrofauna (earthworms, millipedes, ants, termites, and fossorial or earth-dwelling vertebrates) move through macroand micropores in the soil. The macrofauna plays a role in moving parts of the soil profile around, and form many sorts of burrows and pores; they are often termed ‘‘ecological engineers.’’

relationships have several trophic levels, with bacteria and fungi being fed upon by microbe-feeders, such as protozoa, nematodes and microarthropods, which are in turn preyed upon by predatory nematodes or mites, and these in turn fed upon by higher predators (Fig. 1). In forests or no tillage agroecosystems, where fungi dominate in the surface litter, the dominant flows of energy and nutrients will go via fungal pathways. In contrast, in conventional tilled fields where the organic matter is incorporated in the plow layer (usually 6–8 in. ¼ 15–20 cm), the dominant flows of energy and nutrients may be more bacterially dominated, usually decomposing faster than in the no tillage system.[1]

ZONES OF INFLUENCE The heterogeneous distribution of food resources in the soil matrix makes it difficult to sample adequately for abundances and activities of the biota in a repeatable fashion. A useful approach is to consider soils as being comprised of zones of influence (ZOI), that can be targeted for further study. These ZOI, also termed ‘‘hot spots,’’ are located in the root-rhizosphere, in regions of organic detritus accumulation, or detritusphere, and also in earthworm-influenced regions, such as burrows, which are termed a drilosphere[4] (Fig. 2). These ZOI may represent less than 10% of the volume of the surface A horizon, but account for up to 90% of the total biological activity in soils worldwide.

SOIL FOOD WEBS The initial breaking up or ‘‘comminution,’’ of plant litter (above- and below-ground) results from the chewing and macerating action of both large and small animals. This comminution benefits the fauna, which derive nutritional benefit from the litter and= or microbes initially colonizing the plant material. The increased surface area and further inoculation of the smaller pieces enhances the microbial access to, and breakdown of, these tissues. The decomposition process drives complex food webs[2,3] in the soil, with numerous interactions between the initial agents of decomposition, the bacteria and fungi, and the fauna that in turn feed upon them, which facilitates nutrient return in the soil matrix (Fig. 1). These feeding 1222 Copyright © 2006 by Taylor & Francis

ROLES OF BIOTA IN SOIL FOOD WEBS The functional roles of soil organisms can be compared most usefully in terms of body width. The microbes and microfauna inhabit soil water films, and are restricted to this aquatic milieu. In contrast, the mesoand macrofauna, from acari (mites) to earthworms, inhabit gas-filled pores, and move around in the soil matrix for considerable distances (Fig. 3).[5] Bacteria Bacteria are unicellular prokaryotes (organisms lacking a unit membrane-bounded nucleus and other Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001817 Copyright # 2006 by Taylor & Francis. All rights reserved.

Organisms and Soil Food Webs

1223

Fig. 1 Representation of detrital food web in shortgrass prairie. Fungal-feeding mites are separated into two groups (I and II) to distinguish the slow-growing oribatids from faster-growing taxa. Flows omitted from the figure for the sake of clarity include transfers from every organism to the substrate pools (death) and transfers from every animal to the substrate pools (defecation) and to inorganic N (ammonification). (From Ref.[2].)

organelles), that are found in all habitats on earth. They are exceedingly numerous (more than 1030, or one million trillion trillion[6]) and diverse, currently comprising over 35 kingdoms in two domains, the Archaea and Eubacteria. They are active in all aspects of elemental cycling, and needed for nitrogen cycling, both in nitrogen fixation (splitting N2 and incorporating N into organic compounds), and subsequent transformational pathways as well. They are also primary agents of decomposition in many habitats, and are particularly active in rhizospheres.[4]

decomposed subunits, and translocating them back through the hyphal network. Fungi are very abundant, particularly in undisturbed forest floors, in which literally thousands of kilometers of hyphal filaments will occur per gram of leaf litter. The roles of mycorrhizas (literally ‘‘fungus–root,’’ or symbiotic fungi associated with many plants) in soil systems are being increasingly viewed as central to much of terrestrial ecosystem function. Mycorrhizas are essential to the growth and reproduction of numerous families of plants.[1]

Fungi

Microfauna

Fungi are multicellular eukaryotes that are found in many habitats worldwide. They have long, ramifying strands (hyphae) which can grow into and explore many microhabitats, and are used for obtaining water and nutrients. The hyphae secrete a considerable array of enzymes, such as cellulases, and even lignases in some specialized forms (useful in breaking down wood), decomposing substrates in situ, taking up the

The unicellular eukaryotes, or Protoctista, are more often called protozoans. They include the flagellates, naked amoebae, testacea, and ciliates. These organisms range in size from a few cubic micrometers in volume (micro flagellates) to larger ciliates, which may be up to 500 mm in length and 20–30 mm in width. Protozoa are abundant, reaching densities of from 100 to 200 thousand per gram of soil. Bacteria, their principal


1224


Fig. 2 Areas of activity in soil systems. These ‘‘ZOI’’ may be 90% of the total biological activity in most soils worldwide. (From Ref.[4].)

prey, often exist in numbers up to one billion per gram of soil. All of these organisms are true water-film dwellers, and become dormant or inactive during episodes of drying in the soil. They can exist in inactive or resting stages literally for decades at a time in very xeric environments. Mesofauna Nematodes Nematodes have a wide range of feeding preferences. A general trophic grouping is bacterial feeders, fungal feeders, plant feeders, and predators and omnivores. Anterior (stomal or mouth) structures can be used to differentiate general feeding or trophic groups. Because nematodes reflect the developmental stages of the systems in which they occur (e.g., annual vs.


perennial crops, or old fields and pastures and more mature forests), they have been used as indicators of overall ecosystem condition.[7,8] Collembola Collembolans, or ‘‘springtails’’ are primitive apterygote (wingless) insects. They are called ‘‘springtails’’ because many of them have a spring-like lever, or furcula, which enables them to move many body-lengths away from predators in a springing fashion. Collembolans are ubiquitous members of the soil fauna, often reaching abundances on 100,000 or more per m2. They occur throughout the soil profile, where their major diet is decaying vegetation and associated microbes (usually fungi). However, like many members of the soil fauna, collembolans defy placement in exact trophic groups. Many collembolan species will eat


1225

Fig. 3 Size classification of organisms in decomposer food webs by body width. (From Ref.[5].)

nematodes when those are abundant. Some feed on live plants or their roots. One family (Onychiuridae) may feed in the rhizosphere and consume mycorrhizas or even plant pathogenic fungi.[9]

sources and are rare except in agricultural soils. The Prostigmata contains a broad diversity of mites with several feeding habits.[1]

Macrofauna Mites (Acari) The soil mites, Acari, are chelicerate arthropods related to the spiders. They are often the most abundant microarthropods in soils. A 100 g sample may contain as many as 500 mites representing nearly 100 genera. This diverse array includes participants in three or more trophic levels, with varied strategies for feeding, reproduction, and dispersal. The oribatid mites (Oribatei) are the characteristic mites of the soil and are usually fungivorous or detritivorous. Mesostigmatid mites are nearly all predators on other small fauna, although a few species are fungivores and may become numerous at times. Astigmatid mites are associated with rich, decomposing nitrogen


Termites (Isoptera) are one of the major ecosystem ‘‘engineers,’’ particularly in tropical regions. Termites are social insects with a well-developed caste system. By their ability to digest wood, they have become economic pests of major importance in some regions of the world. The termites in a primitive family, the Kalotermitidae, possess a gut flora of protozoans, which enables them to digest cellulose. Their normal food is wood that has come into contact with soil. Many species of termites construct runways of soil, or along root channels, and some are builders of large, spectacular mounds. Members of the phylogenetically advanced family Termitidae possess a formidable array of microbial symbionts (bacteria and fungi, but not protozoa),

1226

and elevated pH in their hindguts,[10] which enable them to process and digest the humified organic matter in tropical soils and to thrive on it. Termites parallel earthworms in ingestive and soil turnover functions. The principal difference is that earthworms egest much of what they ingest in altered form (that enriches microbial action), whereas termites transfer large amounts of soil (organic material into building nests and mounds (carbon sinks).

Earthworms Much of the evidence for earthworm effects on soil processes comes from agroecosystems and involves a small group of European lumbricids (Lumbricidae family in the Oligochaeta order). Impacts of exotic earthworms on native species are not well understood, although there is evidence that when native habitat is destroyed and native earthworm species extirpated, exotic earthworms colonize the newly empty habitat. As more extensive studies are carried out, it is becoming clear that earthworms are present in a wide variety of tropical as well as temperate ecosystems. Earthworms have important roles in the fragmentation, breakdown and incorporation of soil organic matter (SOM). This affects the distribution of SOM, and also its chemical and physical characteristics. Changes in any of these soil parameters may have significant effects on other soil biota, by changing their resource base (e.g., distribution and quality of SOM, microbes or microarthropods) or by changing the physical structure of the soil.[11] Earthworm activities impact the communities of other soil biota through their effects on the chemical and physical characteristics of SOM, causing changes in the microbial and microarthropod communities, and also having impacts elsewhere in the soil food web.[12]

CONCLUSIONS Soil biota are very interconnected with each other by a variety of trophic and nutrient-flow pathways. The possibility of enhancing biotic activity in agricultural systems via conservation or no tillage regimes is very great, and much research is now focusing on this area of interest. This approach permits a melding of



interests between those who want to reduce fossil fuel inputs to agroecosystems and those who are concerned with enhancing carbon sequestration in soils as well.

REFERENCES 1. Coleman, D.C.; Crossley, D.A., Jr. Fundamentals of Soil Ecology; Academic Press: San Diego, CA, 1996; 205. 2. Hunt, H.W.; Coleman, D.C.; Ingham, E.R.; Ingham, R.E.; Elliott, E.T.; Moore, J.C.; Rose, S.L.; Reid, C.P.P.; Morley, C.R. The detrital food web in a shortgrass prairie. Biol. Fert. Soils 1987, 3, 57–68. 3. Moore, J.C.; de Ruiter, P.C. Invertebrates in detrital food webs along gradients of productivity. In Invertebrates as Webmasters in Ecosystems; Coleman, D.C., Hendrix, P.F., Eds.; CABI: Wallingford, UK, 2000; 161–184. 4. Beare, M.H.; Coleman, D.C.; Crossley, D.A., Jr.; Hendrix, P.F.; Odum, E.P. A hierarchical approach to evaluating the significance of soil biodiversity to biogeochemical cycling. Plant Soil 1995, 170, 5–22. 5. Swift, M.J.; Heal, O.W.; Anderson, J.M. Decomposition in Terrestrial Ecosystems; Univ. of California Press: Berkeley, CA, 1979; 379. 6. Whitman, W.B.; Coleman, D.C.; Wiebe, W.J. Prokaryotes: the unseen majority. Proc. Natl. Acad. Sci. 1998, 95, 6578–6583. 7. Bongers, T. The maturity index: an ecological measure of environmental disturbance based on nematode species composition. Oecologia 1990, 83, 14–19. 8. Yeates, G.W.; Bongers, T.; de Goede, R.G.M.; Freckman, D.W.; Georgieva, S.S. Feeding habits in soil nematode families and genera—an outline for soil ecologists. J. Nematol. 1993, 25, 315–331. 9. Lartey, R.T.; Curl, E.A.; Peterson, C.M. Interactions of mycophagous collembola and biological control fungi in the suppression of Rhizoctonia Solani. Soil Biol. Biochem. 1994, 26, 81–88. 10. Bignell, D.E.; Eggleton, P. On the elevated intestinal pH of higher termites (isoptera termitidae). Insectes Sociaux 1995, 42, 57–69. 11. Hendrix, P.F., Ed.; Earthworm Ecology and Biogeography in North America; Lewis Publishers: Boca Raton, FL, 1995; 335. 12. Fu, S.; Cabrera, M.L.; Coleman, D.C.; Kisselle, K.W.; Garrett, C.J.; Hendrix, P.F.; Crossley, D.A., Jr. Soil carbon dynamics of conventional tillage and no-till agroecosystems at Georgia Piedmont—HSB-C models. Ecol. Model. 2000, 131, 229–248.

Organo-Mineral Relationships Claire Chenu Alain F. Plante Unite´ de Science du Sol, INRA-Versailles, Versailles, France

Pascale Puget ESITPA, Rouen, France

INTRODUCTION Organo-mineral relationships are a fundamental feature of soils and are often used to define and differentiate soils from geological parent materials. Soil is primarily a mineral matrix (except in organic soils such as bog peat), but receives inputs of organic materials from various natural sources such as litterfall, root exudates, or from various anthropogenic sources such as manure additions. Organo-mineral relationships range in degree of association from the spatial distribution of particulate organic matter and mineral particles with minimal interaction to the inseparable organic matter intercalated between clay layers. Here, the term association is used to describe an arrangement of organic matter and minerals that has an undefined degree of cohesion. The term complex is restricted to cases where adsorption is the dominant mechanism. Organo-mineral associations cover a wide range of spatial scales, from nanometric (e.g., the complex of an organic acid with a clay sheet) to decimetric (e.g., a soil ped); however, they all originate at the molecular scale where physico-chemical interactions lead to the formation of bonds. The formation of organo-mineral associations may require only minutes, but their lifetime is related to that of the organic component, which can be short or extend to thousands of years. The impact of organo-mineral complexes in soils is significant; 40–80% of soil carbon is present in clay-sized separates and cannot be separated easily from clay minerals.[1] Their importance is also qualitative and functional because the formation of organo-mineral complexes in soils strongly impacts the properties of the mineral phase and influences the biodegradation of organic matter.

CHARACTERIZATION OF ORGANO-MINERAL RELATIONSHIPS Historically, organic matter has been studied by methods that involved their solubilization and separation from soil minerals. On the other hand, mineralogists Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006622 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

analyzed soil minerals after comprehensive destruction of organic matter with H2O2. Therefore, most of the past knowledge concerning organo-mineral associations was gained from studies performed on synthetic organomineral complexes, prepared from well-defined constituents such as clay minerals from geological deposits and pure organic compounds (e.g., polysaccharides, proteins).[2] Physical methods for soil fractionation such as size or density separation now enable the separation of intact organo-mineral associations from soils and their direct study with a wide range of nondestructive methods such as nuclear magnetic resonance spectroscopy.[3] Primary organo-mineral associations are associations of organic matter with individual soil mineral particles.[4] Their separation requires complete dispersion of soil using mechanical means and size or density separation techniques. Secondary organo-mineral associations, namely aggregates, are made by grouping together of several primary organo-mineral associations. Secondary organo-mineral associations are normally isolated from soil by sieving techniques after mild soil dispersion (e.g., a simple immersion of soil sample in water and agitation).[5] In primary organo-mineral associations, more organic matter is more closely associated with minerals as particle size decreases: sand-sized organic matter (i.e., decomposing plant debris) is separated easily from sand-sized minerals by densimetric techniques,[6,7] whereas little free mineral or organic matter seem to occur in the clay-sized fraction.[8] A variety of compounds are involved in primary organo-mineral associations. The clay-sized associations consist of highly processed and humified compounds, as well as microorganisms and their metabolites.[1,4,9] Clay– organic matter associations also show diverse microstructures, from organic interlayer complexes[10,11] and organic coatings on clay minerals surfaces to more complex structures such as microaggregates of bacteria or plant debris with clay particles (Fig. 1).[12,13] In various soil types and management contexts, secondary organo-mineral complexes are hierarchical: particles associate into microaggregates, which in turn 1227

1228

Fig. 1 Microstructure of primary organo-mineral associations observed with transmission electron microscopy in clay-sized fractions from soils: (A) complex microaggregate, (B) microaggregate in which OM occurs as layers between stacks of clay. (From Ref.[57].)

associate into larger aggregates.[14–17] The smaller the aggregate hierarchical level, the higher is its physical stability, because the binding agents involved change with the spatial scale. Clay-sized associations are bound by sesquioxides, humic materials, and polysaccharides. These associations are bound into 250 mm by transient agents (e.g., polysaccharides) and temporary ones (e.g., fine roots and fungal hyphae) (Fig. 2).[18,19] Aggregates appear to be formed and stabilized around decomposing plant debris that act as hot spots for microbial activity (Fig. 3).[16,20–23] Binding Forces Within Organo-Mineral Associations A variety of bonds may be formed when organic molecules in solution come into contact with the surface of


Organo-Mineral Relationships

Fig. 2 Nature and scale of various organo-mineral associations. (From Ref.[18].)

soil minerals (see Fig. 4).[24] For small organic molecules, these interactions depend on their ability to establish electrostatic interactions or dipole interactions. Polarity and polarizability are thus important adsorption-related properties of such organic molecules. For macromolecules, weak interactions such as van der Waals interactions and hydrogen bonding become very important. The conformation of macromolecules largely influences their adsorption, which increases with molecular weight.[2] High-molecularweight polymers are generally adsorbed in an irreversible fashion and therefore the adsorption of macromolecules into soil minerals often leads to the formation of stable organo-mineral complexes. Organic polymers such as polysaccharides, proteins, fulvic, and humic acids are active in the formation of organo-mineral complexes because of their high molecular weight and charge. Similarly, clay minerals are the most active mineral constituents in the formation of organo-mineral complexes because of their charge and their high specific


1229

contacts between the organic and mineral surfaces and thus aiding the creation of new bonds.[32,33]

ORGANO-MINERAL RELATIONSHIPS ALTER THE PROPERTIES OF SOIL MINERALS

Fig. 3 Schematic presentation of the interaction between soil organic matter decomposition and aggregate formation and destruction. (Adapted from Refs.[21,22,56].)

surface area.[2] Solid particles, such as bacteria, fungi, or plant debris also establish physico-chemical interactions with mineral surfaces through the process of adhesion. Physico-chemical interactions at the molecular scale (i.e., adsorption and adhesion) provide tensile forces that increase both the tensile strength and the compressive strength of primary as well as secondary organomineral complexes and associations.[25–27] At a larger scale, the binding of aggregates also involves physical mechanisms. Fungal hyphae and fine roots stabilize aggregates by entangling the soil particles[28–30] and thereby increasing aggregate strength through compressive forces. Physical processes, such as those related to shrink–swell or freeze–thaw also alter organomineral associations by (i) creating failure zones that separate aggregates,[14,31] or by (ii) increasing the

Soil properties observed at a macroscopic scale are largely due to changes in the properties and associations of soil primary particles at a microscopic scale, in particular that of clay minerals. For example, the basic physical and physico-chemical mechanisms by which organic matter stabilizes soil aggregates against the disruptive action of water are an increased cohesion of the aggregate or a decreased wettability. Increased cohesion helps the aggregate withstand mechanical disruption by raindrops and resist forces exerted by compressed air upon wetting. Adsorption of large and flexible organic polymers increases the tensile strength of minerals (see above). Decreased wettability of aggregate surfaces slows the rate of wetting of the aggregates and thus the extent of slaking. Synthetic complexes of clays and humic substances, as well as natural complexes, are less wettable than pure clays.[34–36] Associated organic matter also changes the swelling of clays and organo-mineral associations.[37–39] Organo-mineral associations alter the reactivity of soil minerals, particularly of soil clays, thereby altering the retention of cations, trace metals, or organic pollutants in soils. Associated organic matter increases the cation-exchange capacity of soils and soil clays,[40–42] which contributes to the preferential retention of heavy metals by clay-sized fractions of soils. Increased hydrophobicity in the clay fraction, caused by associated organic matter, facilitates the sorption of nonpolar organic pollutants.[43]

Fig. 4 Mechanisms of interaction between clay minerals and organic molecules (From Paul and Clark, 1996).


1230

ORGANO-MINERAL RELATIONSHIPS REDUCE THE BIODEGRADATION RATE OF ORGANIC MATTER Many easily decomposed compounds are retained in soil for longer periods than expected from their biochemistry. Association of organic matter with soil mineral particles and aggregates usually results in a decrease in the rate of biodegradation of the organic matter. Bartlett and Doner[44] measured the decomposition of two radiolabeled-amino acids in four treatments: in a nonaggregated and nonadsorbed state; in a nonaggregated, but adsorbed state; in an aggregated, but nonadsorbed state; and in an aggregated and adsorbed state. The results showed that decomposition was more rapid in the nonaggregated versus aggregated state, and that adsorption further decreased decomposition. The experiment clearly demonstrates two major mechanisms involved in organic-matter protection: ‘‘chemical protection’’ and ‘‘physical protection.’’[45] Chemical protection refers to the reduced availability to microorganisms (and therefore decomposition) of organic matter in soil due to chemical interactions with soil constituents, such as complexation or adsorption on reactive mineral surfaces. Many natural and contaminant organic compounds are not available to microorganisms when adsorbed.[46–49] Models that best describe the decomposition of compounds in soil couple biodegradation with sorption, and assume that adsorbed compounds are not degraded.[50] While some organisms can access some adsorbed compounds, the rate of decomposition is largely controlled by the desorption of the compound.[50,51] The effect of minerals is also indirect because extracellular enzymes adsorb to clay minerals and become inactivated.[52] The physical protection of organic-matter is the reduction in organic-matter decomposition rates caused by the architecture of the organo-mineral soil matrix. Evidence of physical protection comes primarily from experiments or agronomic situations in which soil structure is disrupted, leading to a flush of mineralization.[53–56] A large proportion of labile soil organic-matter appears to be physically protected in microaggregates.[53] Some degree of contact between a microorganism or extracellular enzyme and the organic substrate is needed to allow biodegradation. Soil microorganisms generally occupy less than 2% of soil surface area and less than 1% of soil porosity.[57] Therefore, large distances can exist between organism and substrate. Soil structure controls the continuity of pores filled with water and their accessibility to microorganisms. Habitable pore space consists of 15–25% of the total porosity, while the remainder is inaccessible to microorganisms because pore necks are smaller that 0.2 mm or pores are air-filled rather



than water-filled.[58] Furthermore, soil structure may create sites in which the local physico-chemical conditions are unfavorable to mineralization due to oxygen depletion in anoxic center of aggregates.[59] While the chemical protection provided by adsorption is often permanent, physical protection is highly dynamic. Soil structure is continuously changing due to the action of soil fauna, root growth, wet–dry cycles, and tillage (see Fig. 3). Therefore, physical protection depends on the life expectancy of the arrangement of the protection sites, and is therefore subject to changes due to soil aggregate turnover.[60] While the relative contribution of physical protection to the stabilization of organic matter may be low, it often affords the time for more permanent chemical-protection mechanisms to occur. CONCLUSIONS The existence of ‘‘primary particles’’ and ‘‘clean’’ mineral surfaces in natural soils occurs only rarely (except in the case of sands). Instead, soil is an intimate association of organic matter and mineral surfaces that interact to various degrees. While the binding of organic matter with mineral surfaces occurs at the nano- to micro-meter scale, its impact is felt upwards to the centimeter and meter scales. Organo-mineral relationships are dynamic processes that alter the properties of both the organic matter and mineral particles involved.

REFERENCES 1. Christensen, B.T. Carbon in primary and secondary organo-mineral complexes. In Structure and Organic Matter Storage in Agricultural Soils; Carter, M.R., Stewart, B.A., Eds.; CRC Press Inc.: Boca Raton, FL, 1995; 97–165. 2. Theng, B.K.W. Formation and Properties of Clay– Polymer Complexes; Elsevier Scientific Publishing Company: Amsterdam, 1979. 3. Ko¨gel-Knabner, I. Analytical approaches for characterizing soil organic matter. Org. Geochem. 2000, 31 (7=8), 609–625. 4. Christensen, B.T. Physical fractionation of soil and organic matter in primary particle size and density separates. In Advances in Soil Science; Stewart, B.A., Ed.; Springer: New York, 1992; Vol. 20, 1–90. 5. Elliott, E.T. Aggregate structure and carbon, nitrogen and phosphorus in native and cultivated soils. Soil Sci. Soc. Am. J. 1986, 50 (3), 627–633. 6. Feller, C. Une me´thode de fractionnement granulome´trique de la matie`re organique des sols. Application aux sols tropicaux a` textures grossie`res, tre`s pauvres en humus. Cah. ORSTOM, Se´r. Pe´dologie 1979, XVII, 339–345.


7. Balesdent, J.; Pe´traud, J.P.; Feller, C. Effet des ultrasons sur la distribution granulome´trique des matie`res organiques des sols. Science Sol. 1991, 29 (2), 95–106. 8. Turchenek, J.M.; Oades, J.M. Fractionation of organomineral complexes by sedimentation and density techniques. Geoderma 1979, 21 (4), 311–343. 9. Oades, J.M. An introduction to organic matter in mineral soils. In Minerals in Soil Environments; Dixon, D.E., Ed.; SSSA Book Series No. 1; Soil Science Society of America: Madison, WI, 1989; 89–158. 10. Theng, B.K.G.; Chuchman, G.J.; Newman, R.H. The occurence of interlayer clay–organic complexes in two New Zealand soils. Soil Sci. 1986, 142 (5), 262–266. 11. Righi, D.; Dinel, H.; Schulten, H.R.; Schnitzer, M. Characterization of clay–organic matter complexes resistant to oxidation by hydrogen peroxide. Eur. J. Soil Sci. 1995, 46 (3), 423–429. 12. Feller, C.; François, C.; Villemin, G.; Portal, J.M.; Toutain, F.; Morel, J.L. Nature des matie`res organiques associeés aux fractions argileuses d’un sol ferrallitique. C.R. Acad. Sci. Paris. 1991, 312 (12), 1491–1497. 13. Chenu, C.; Arias, M.; Besnard, E. The influence of cultivation on the composition and properties of clay– organic matter associations in soils. In Sustainable Management of Organic Matter; Rees, R.M., Ball, B.C., Campbell, C.D., Watson, C.A., Eds.; CABI International: Wallingford, UK, 2001; 207–213. 14. Dexter, A.R. Advances in characterization of soil structure. Soil Tillage Res. 1988, 11 (3=4), 199–238. 15. Oades, J.M.; Waters, A.G. Aggregate hierarchy in soils. Aust. J. Soil Res. 1991, 29 (6), 815–828. 16. Puget, P.; Chenu, C.; Balesdent, J. Dynamics of soil organic matter associated with primary particle size fractions of water-stable aggregates. Eur. J. Soil Sci. 2000, 51 (4), 595–605. 17. Six, J.; Paustian, K.; Elliott, E.T.; Combrink, C. Soil structure and organic matter: I. Distribution of aggregate-size classes and aggregate-associated carbon. Soil Sci. Soc. Am. J. 2000, 64 (2), 681–689. 18. Tisdall, J.M.; Oades, J.M. Organic matter and waterstable aggregates. J. Soil Sci. 1982, 33 (2), 141–163. 19. Jastrow, J.D.; Miller, R.M.; Lussenhop, J. Contributions of interacting biological mechanisms to soil aggregate stabilization in restored prairie. Soil Biol. Biochem. 1998, 30 (7), 905–916. 20. Beare, M.H.; Hendrix, P.F.; Coleman, D.C. Waterstable aggregates and organic matter fractions in conventional-tillage and no-tillage soils. Soil Sci. Soc. Am. J. 1994, 58 (3), 777–786. 21. Golchin, A.; Oades, J.M.; Skjemstad, J.O.; Clarke, P. Soil structure and carbon cycling. Aust. J. Soil Res. 1994, 32 (5), 1043–1068. 22. Chenu, C.; Puget, P.; Balesdent, J. Clay–Organic Matter Associations in Soils: Microstructure and Contribution to Soil Physical Stability, 16th World Congress of Soil Science, Montpellier, France, Aug 20–26, 1998; Cirad: Montpellier, France, CD-ROM. 23. Golchin, A.; Baldock, J.A.; Oades, J.M. A model linking organic matter decomposition, chemistry, and aggregate dynamics. In Soil Processes and the Carbon Cycle;


1231

24.

25.

26.

27.

28. 29.

30.

31.

32.

33.

34.

35.

36.

37.

38.

39. 40.

Lal, R., Kimble, J.M., Follett, R.F., Stewart, B.A., Eds.; CRC Press Inc.: Boca Raton, FL, 1998; 245–266. Mortland, M.M. Clay–organic matter complexes and interactions. In Advances in Agronomy; Bray, N.C., Ed.; Academic Press: New York, 1970; Vol. 22, 75–117. Dowdy, R.H. The effect of organic polymers and hydrous oxides on the tensile strength of clay. In Soil Conditioners; Stewart, B.A., Ed.; SSSA: Madison, WI, 1975; 25–33. Chenu, C.; Gue´rif, J. Mechanical strength of clay minerals as influenced by an adsorbed polysaccharide. Soil Sci. Soc. Am. J. 1991, 55 (4), 1076–1080. Czarnes, S.; Hallett, P.D.; Bengough, A.G.; Young, I.M. Root- and microbial-derived mucilages affect soil structure and water transport. Eur. J. Soil Sci. 2000, 51 (3), 435–443. Tisdall, J.M. Fungal hyphae and structural stability of soil. Aust. J. Soil Res. 1991, 29 (6), 729–743. Dorioz, J.M.; Robert, M.; Chenu, C. The role of roots, fungi and bacteria on clay particle organization: an experimental approach. Geoderma 1993, 56 (1–4), 179–194. Degens, B.P. Macro-aggregation of soils by biological bonding and binding mechanisms and the factors affecting these: a review. Aust. J. Soil Res. 1997, 35 (3), 431–459. Preston, S.; Griffiths, B.S.; Young, I.M. Links between substrate additions, native microbes, and the structural complexity and stability of soils. Soil Biol. Biochem. 1999, 31 (11), 1541–1547. Haynes, R.J.; Swift, R.S. Stability of soil aggregates in relation to organic constituents and soil water content. J. Soil Sci. 1990, 41 (1), 73–83. Caron, J.; Kay, B.D.; Stone, J.A. Improvement of aggregate stability of a clay loam with drying. Soil Sci. Soc. Am. J. 1992, 56 (5), 1583–1590. Jouany, C. Surface free energy components of clay–synthetic humic acid complexes from contact-angle measurements. Clays Clay Mines. 1991, 39 (1), 43–49. Jan˜czuk, B.; Hajnos, M.; Bialopiotrowicz, T.; Kliszcs; Bilin˜ski, B. Hydrophobization of the soils by dodecylammonium hydrochloride and changes of the components of its surface energy. Soil Sci. 1990, 150 (5), 792–798. Chenu, C.; Bissonnais, Y.l.; Arrouays, D. Organic matter influence on clay wettability and soil aggregate stability. Soil Sci. Soc. Am. J. 2000, 64 (4), 1479–1486. Chenu, C. Influence of a fungal polysaccharide, scleroglucan, on clay microstructures. Soil Biol. Biochem. 1989, 21 (2), 299–305. Chenu, C. Clay- or sand-polysaccharides associations as models for the interface between microorganisms and soil: water-related properties and microstructure. Geoderma 1993, 56 (1–4), 143–156. Emerson, W.W. Water retention, organic C and soil texture. Aust. J. Soil Res. 1995, 33 (2), 241–351. Thompson, M.L.; Zhang, H.; Kazemi, M.; Sandor, J. Contribution of organic matter to cation exchange capacity and specific surface area of fractionated soil materials. Soil Sci. 1989, 148 (4), 250–257.

1232

41. Curtin, D.; Rostad, H.P.W. Cation exchange capacity and buffer potential of saskatchewan soils estimated from texture, organic matter and pH. Can. J. Soil Sci. 1997, 77 (4), 621–626. 42. Bigorre, F. Contribution des argiles et des matie`res organiques a` la re´tention d’eau dans les Sols. Significa´ change en tion et ro ^ le fondamental de la capacite´ d’E cations. C.R. Acad. Sci. Paris 1999, 330 (3), 245–250. 43. Wershaw, R.L. A new model for humic materials and their interactions with hydrophobic organic chemicals in soil water or sediment water systems. J. Contaminant Hydrol. 1986, 1 (1/2), 29–45. 44. Bartlett, J.R.; Doner, H.E. Decomposition of lysine and leucine in soil aggregates: adsorption and compartmentalization. Soil Biol. Biochem. 1988, 20 (5), 755–759. 45. Baldock, J.A.; Skjemstad, J.O. Role of the soil matrix and minerals in protecting natural organic materials against biological attack. Org. Geochem. 2000, 31 (7/8), 697–710. 46. Aardema, B.W.; Lorenz, M.G.; Krumbein, W.E. Protection of sediment adsorbed transforming DNA against enzymatic inactivation. Appl. Environ. Microbiol. 1983, 46 (2), 417–420. 47. Ogram, A.; Jessup, R.E.; Ou, L.T.; Rao, P.S.C. Effects of sorption on biological degradation rates of (2,4Dichlorophenoxy) acetic acid in soils. Appl. Environ. Microbiol. 1985, 49 (3), 582–587. 48. Dashman, T.; Stotzky, G. Microbial utilization of amino acids and a peptide bound on homoionic montmorillonite and kaolinite. Soil Biol. Biochem. 1986, 18 (1), 5–14. 49. Jones, D.L.; Edwards, A.C. Influence of sorption on the biological utilization of two simple carbon substrates. Soil Biol. Biochem. 1998, 30 (14), 1895–1902. 50. Scow, K.M.; Johnson, C.R. Effect of sorption on biodegradation of soil pollutants. Adv. Agron. 1997, 58 (1), 1–56. 51. Rao, P.S.C.; Bellin, C.A.; Lee, L.S. Sorption and biodegradation of organic contaminants in soils: conceptual representations of process coupling. In Environmental Impact of Soil Component Interactions: Volume 1:



52.

53.

54.

55.

56.

57.

58.

59.

60.

Natural and Anthropogenic Organics. Proceedings of a Workshop Entitled ‘Impact of Interactions of Inorganic, Organic and Microbiological Soil Components on Environmental Quality, Edmonton AB, Aug 11–15, 1992; Lewis Publishers: Boca Raton, FL, 1995; 263–274. Quiquampoix, H. Mechanisms of protein adsorption on surfaces and consequences for soil extracellular enzyme activity. In Soil Biochemistry; Bollag, J.M., Stotzky, G., Eds.; Marcel Dekker: New York, 2000; Vol. 10, 171–206. Gregorich, E.G.; Kachanovski, R.G.; Voroney, R.P. Carbon mineralization in soil size fractions after various amounts of aggregate disruption. J. Soil Sci. 1989, 40 (3), 649–659. Balesdent, J.; Boisgontier, D.; Mariotti, A. Effect of tillage on soil organic carbon mineralization estimated from 13C abundance in maize fields. J. Soil Sci. 1990, 41 (4), 587–596. Beare, M.H.; Cabrera, M.L.; Hendrix, P.F.; Coleman, D.C. Aggregate-protected and unprotected organic matter pools in conventional-tillage and no-tillage soils. Soil Sci. Soc. Am. J. 1994, 58 (3), 787–795. Balesdent, J.; Chenu, C.; Balabane, M. Relationship of soil organic matter dynamics to physical protection and tillage. Soil Tillage Res. 2000, 53 (3/4), 215–230. Chenu, C.; Stotzky, G. Interactions between microorganisms and soil particles: an overview. In Interactions Between Microorganisms and Soild Particles in the Soil Environment; Huang, P.M., Ed.; 2002, in press. Hassink, J.; Bouwman, J.; Brussaard, L.B. Relationships between habitable pore space, soil biota and mineralization rates in grassland soils. Soil Biol. Biochem. 1993, 25 (1), 47–55. Sextone, A.J.; Revsbech, N.P.; Parkin, T.B.; Tiedje, J.M. Direct measurement of oxygen profiles and denitrification rates in soil aggregates. Soil Sci. Soc. Am. J. 1985, 49 (3), 645–651. Plante, A.F. Soil aggregate turnover and the physical protection of soil organic matter as measured using dy-labelled tracer spheres. Ph.D. thesis, University of Alberta, Edmonton, Canada, AB, 2001.

Oxygen Diffusion Rate and Plant Growth Witold Ste¸ pniewski Technical University of Lublin and Polish Academy of Sciences, Lublin, Poland

INTRODUCTION

SOIL PARAMETERS AND ODR

What Is Oxygen Diffusion Rate?

The ODR index reflects comprehensively the soiloxygen availability, as it comprises the effect of all the factors such as respiration, moisture, and the physical particle arrangement that influence the concentration of oxygen in soil air, the effective thickness of the water films around the roots, and their diffusion characteristics. The ODR values in soils vary, most frequently, within a range from 0 to 200 mg m2 sec1 (Figs. 1–3). They decrease with soil-moisture content and with soil compaction and they increase with moisture tension, with air-filled porosity of the soil, and with oxygen content in soil air.[4] Due to this, a decrease in ODR is expected usually with depth, especially just above the ground water table. An example of the relationship of ODR on soil-moisture tension and bulk density is presented in Fig. 1. Beyond the soil moisture content and bulk density, the value of ODR is related to many other soil parameters. The dependence on the gas diffusion coefficient[6] is shown in Fig. 2. The relationship with air permeability of the soil[6] is illustrated in Fig. 3. The value of ODR is related also to redox potential[4] and to dehydrogenase activity.[7]

Oxygen diffusion rate (ODR) is an electrochemical method of assessment of soil oxygen availability to plant roots. It is based on the analogy of oxygen uptake by plant roots and by platinum wire electrode placed in the soil. Oxygen diffusion rate is of importance because availability of oxygen to plant roots is a basic factor of soil productivity. A knowledge of the method of measurement of ODR helps one to evaluate oxygen requirements of particular plant species at different stages of their development. Moreover, we can use it as a diagnostic tool to assess the oxygen availability in a particular soil under definite conditions. Concept and Principle of the Method As early as 1926, Hutchins[1] expressed a conviction that it is not so much the concentration of oxygen in the soil as the possibility of its uptake by plant roots that determines the plant response to the oxygen conditions in the soil. This idea was realised in 1952 by Lemon and Erickson,[2] who designed an electrode simulating oxygen uptake by plant roots. An important role in this concept is played by the presence of water or exudate films on the root surface. These films, due to the low diffusion coefficient of gases in water, form a significant obstacle in the path of oxygen from the soil air to the root.[2,3] The principle of this method consists in the measurement of the amount of oxygen diffusing onto the surface of a platinum wire electrode (usually 0.5–1.0 mm in diameter and several mm in length), where it is reduced electrochemically to hydroxide ions or water.[2,3] In practice, the value of ODR is determined on the basis of measurement of the diffusion current intensity on a platinum electrode, which is negatively polarized in order to provide specific conditions for the reduction of oxygen molecules only. The platinum wire is therefore a model of oxygen absorbing root, and the intensity of oxygen flux to the electrode indicates the maximum amount of oxygen that would be available for a plant root placed in the same spot as the electrode. 1236 Copyright © 2006 by Taylor & Francis

ODR AND PLANT RESPONSE Seedling Emergence In a typical relationship between plant emergence and ODR[4] starting with the highest values of ODR, initially we observe a plateau range with insignificant differentiation in emergence. Then, after reaching a certain limiting value a significant decrease in the final emergence percentage in comparison with the germination capacity is observed. Thereafter, there is a rapid linear decline in the number of emerging plants as the ODR decreases; this number falling to zero at the critical ODR. It was found that the limiting values for the plant emergence range from 25 to 100 mg m2 sec1 and the critical values from 7 to 40 mg m2sec1.[4] Root Response Root tissue is that part of a plant which is subjected first to oxygen deficiency or anoxia, and the Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120006603 Copyright # 2006 by Taylor & Francis. All rights reserved.

Oxygen Diffusion Rate and Plant Growth

1237

Fig. 3 Oxygen diffusion rate vs. air permeability (k) of the soil. Collective data for 24 soil horizons of six different soil profiles from Hungary. (Adapted from Ref.[6].) Fig. 1 Oxygen diffusion rate vs. bulk density and soil moisture tension in a loamy textured chernozem rendzina soil. (Adapted from Ref.[5].)

consequence is a decrease of the root biomass at low ODR.[4,8] It was confirmed that wheat-root population in soil[9] increased linearly with ODR in the interval from 40 to 60 mg m2 sec1 more than 5 fold. Rootelongation rate of three desert shrubs also increased with ODR in the interval 30–90 mg m2 sec1.[10] It was found that for numerous grass species critical ODR value is below 10 mg m2 sec1.[4] In case of lowland rice, root porosity decreased from 23% at

Fig. 2 Oxygen diffusion rate vs. relative gas diffusion coefficient D/D0 in soil, where gas diffusion coefficient D of a gas relates to the soil and D0 to the atmospheric air at the same temperature and pressure conditions. Collective data for 24 soil horizons of six different soil profiles from Hungary. (Adapted from Ref.[6].)


ODR 100 mg m2sec1.[4] Shoot Response Shoot response to root hypoxia expressed by the ODR values has been studied intensively. As it was found, the uptake of water by orange seedling was 2–3 times higher at ODR >60 mg m2 sec1 as compared to that at ODR 5000 m in altitude) emerging from these chains; and 2) nonvolcanic rocks or hard lava flows with notched summits and steep slopes.[4]

SOILS Pa´ramos soils have been described by Jenny since 1948.[6] Soils evolve in function of the convergent effects of low temperature, high soil moisture, and Al availability.[7] In the supra-Pa´ramo, organic matter production decreases with altitude, and soils are very shallow, with umbric epipedon[8] and periglaciar features. In nonvolcanic Venezuelan areas, Cryepts are dominant. The presence of giant rosette plants gives the soil pattern discontinuous properties with higher carbon content only at soil surface and around the rosette.[9] In volcanic areas, during glacier periods, ice caps protected soils from ash deposits. The most recent deposits are weakly weathered and form Vitrandic Cryents. At lower altitudes, in humid conditions, the availability of Al is the second important factor in soil formation. When Al availability is very low, Humods and Fragiaquods are observed on gneiss and micaschists rich in quartz (Podocarpus Paramo in the south of Ecuador). With higher Al availability, Inceptisols (Dystrudept) form with an umbric epipedon. With


1241

high Al availability, Al forms very stable complexes with organic matter.[10] The main pedological process is an acidic complexolitic andosolization with Al þ 1=2 Fe oxalate extract >2% and Al% pyropyrophosphate=Al% oxalate >0.5 characteristic of nonallophanic Andisols,[11] which are the more extended soil types in the Pa´ramos. On recent volcanic ashes, important Al availability leads to a very fast soil formation, with Vitrands forming in less than 1000 years. The farther the deposits are from their emission source, the finer are the ashes, and the faster is the rate of weathering. The older, most evolved soil profiles ( 3000 years) have the lowest base reserve and can been classified as Udands. The nonallophanic andosolization process could also be active on nonvolcanic Al-rich substrates (paleo-oxisols or metamorphic rocks). All of these soils are acidic and have anionic retention capacities (especially for P and S). Pa´ramo Soils and Carbon Content All Pa´ramo soils are very rich in organic carbon. In the Paramo of Northern Peru, soils contain over 10 kg m2 C, whatever the nature of the parent rock (limestones, volcanic rocks, or sandstones).[12] Low temperatures and stable organo-metallic complexes[7] reduce the biological activity and therefore decrease the rate of organic matter mineralization. High carbon accumulation results from different pedologic processes on successive ash layers, but also on colluvial buried soils on steep slopes. Carbon sequestration can exceed 85 kg m2 in the polygenic nonallophanic matured Andisols of Northern Ecuador (Table 2, Photo 1).[13] Altitude increases the soil organic carbon content with a maximum density (reaching 10 g cm3) at around 3900 m. These Andisols have black thick melanic epipedon attributed to the presence of humic acids produced by the degradation of Poaceae family plants,[11] and can been classified as Melanudands (Photo 1).[8] Pa´ramos and Water Cycle Regulation The water content in nonallophanic Andisols at 1500 kPa is over 1000 g kg1 (Hydric properties).[8] Table 2 Carbon Pool (kg m2) of some Pa´ramos soils Place

Pichincha Carchi Cajas Fierro Cerro Sabanilla (1) (2) (3) Urcu (4) Toledo (5) (6)

A

35.6

46.3

B

56.7

86.4

36.4

34.1

15.8

11.2

1: Thaptic Hapludand; 2: Hydric pachic Melanudand; 3: Hydric Melanudand; 4: Humic Andic Hapludox; 5: Humic Dystrudept; 6: Typic Fragiaquod. A: First meter of the profile; B: whole profile.

1242

and can be 2000 g kg1 at field capacity. The formation of an organo-metallic complex network leads to important microporosity. The more mature the soils, the richer they are in humic substances, and the higher is their microporosity. For matured Andisols, micropores with a radius 2=3

c=b < 2=3

Class II

Spheroid

b=a > 2=3

c=b > 2=3

Class III

Blade

b=a < 2=3

c=b < 2=3

Class IV

Rod

b=a < 2=3

c=b > 2=3

a, b, c are the crystallographic axes. (Adapted from Ref.[7].)

Table 2 Roundness grades Class limits Grade term

Ref.

[9]

Very angular

[10]

Ref.

Ref.[11] Rho scalea

0.12–0.17

0.00–1.00

0–0.15

0.17–0.25

1.00–2.00

Subangular

0.15–0.25

0.25–0.35

2.00–3.00

Subrounded

0.25–0.40

0.35–0.49

3.00–4.00

Rounded

0.40–0.60

0.49–0.70

4.00–5.00

Well-rounded

0.60–1.00

0.70–1.00

5.00–6.00

Angular

a

Based on Ref.[10] class limits. (Adapted from Ref.[3].)

These processes are common in diagenesis and weathering processes associated with sedimentary rocks and sedimentary bodies (alluvium and colluvium), and are integral in the rock cycle of igneous, sedimentary, and metamorphic rocks. In soils, shape analysis is used to determine uniformity of parent material and origin and mode of soil formation, and used in mineral and structure analysis. Soil-forming processes at or near the earth’s surface can then be inferred to link parent material origin and genesis of soil materials and soil profiles.

CONCLUSIONS Particle shape classes and secondary clay mineral content have been used in combination to determine soil


texture grades. Texture grades and their arrangement in soil structure largely determine pore volume in soils, which can be related to water and air-filled porosity as well as soil strength which has implications for soil– plant root relationships. Most of the use of particle shape analysis has been confined to soil genesis studies. These studies together with secondary clay mineral formation include interpretations of initial parent materials, inferred factors, and processes used to describe the nature and genesis of soil particulate mineralogical properties and hence, the distribution of contiguous soils in a landscape.

REFERENCES 1. American geological institute. Dictionary of Geological Terms, 2nd Ed.; Dolphin Reference Books: U.S.A., 1962. 2. Pettijohn, F.J.; Potter, P.E.; Siever, R. Sand and Sandstone; Springer: New York, 1987. 3. Brewer, R. Fabric and Mineral Analysis of Soils; Wiley: U.S.A., 1964. 4. Wadell, H. Volume, shape, and roundness of rock particles. J. Geol. 1932, 40, 443–451. 5. Wadell, H. Sphericity and roundness of rock particles. J. Geol. 1933, 41, 310–331. 6. Wadell, H. Volume, shape and roundness of quartz particles. J. Geol. 1935, 43, 250–279. 7. Zingg, Th. Beitragzur Schotteranalyse. Schweiz. Mineral. Petrog. Mitt. 1935, 15, 39–140. 8. Krumbein, W.C. Measurement and geological significance of shape and roundness of sedimentary particles. J. Sediment. Petrol. 1941, 11, 64–72. 9. Pettijohn, F.H. Sedimentary Rocks, 2nd Ed.; Harper Brothers: New York, 1957. 10. Powers, M. A New roundness scale for sedimentary particles. J. Sediment. Petrol 1953, 23, 117–119. 11. Folk, R.L. Student operated error in determination of roundness, sphericity and grain size. J. Sediment. Petrol. 1955, 25, 297–301.

Pedogenic Silica Accumulation Katherine J. Kendrick United States Geological Survey, Pasadena, California, U.S.A.

INTRODUCTION Pedogenic silica is defined as silica that has precipitated in the soil environment, taking the form of opal-A, opal-CT, or microcrystalline quartz. Opaline silica, which is amorphous to poorly crystalline, is the most common form of pedogenic silica.[1,2] Pedogenic silica is formed in arid, semiarid, and Mediterranean climates where there is adequate precipitation to mobilize the silica, but not so much as to leach the silica out of the profile. Pedogenic silica accumulation has been reported throughout the western U.S., Australia, Italy, South Africa, and New Zealand.[3] In its most advanced form, it cements the soil fabric to form extremely hard horizons known as duripans or silcrete.

SOURCES OF PEDOGENIC SILICA Pedogenic silica is derived from two main sources. The weathering and hydrolysis of primary silicate minerals provide for the slow release of silica into solution as silicic acid. Easily weathered silicates, such as olivine and Ca-plagioclase, are particularly likely to contribute silica into solution. The second source is from the weathering of primary amorphous silica, including volcanic glass and biogenic silica (e.g., phytoliths, diatoms). Such amorphous silica is the most soluble form of silica found in the soil environment. Weathering of glassy volcanic products, including ash and tephra, to release silicic acid into solution occurs readily in soils.

DISSOLUTION OF SILICA Factors controlling the dissolution of silica include soil solution pH, the presence of organic matter, the particle size of the material, and the presence of coatings on grains. The solubility of silica is moderate throughout the pH range of most soils, but is particularly high at pH values 9, as found in sodium carbonate soil systems. Organic matter enhances the dissolution of silica, including quartz, and impedes precipitation.[3] The dissolution of quartz has been shown to be highest in the root zone due to organic complexation of monosilicic acid.[4] On the other hand, organic matter may actually form a coating on opal, impeding its dissolution.[5] Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120001755 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

Because of surface area effects, smaller particles are more susceptible to dissolution. This phenomenon accounts for the fact that quartz does not persist in the clay fraction of most soils.[6] Aluminum- and Fe-oxyhydroxides can form insoluble coatings on grains, thereby decreasing rates of dissolution from those grains.[3]

MOVEMENT AND PRECIPITATION OF SILICA Silica is moved in solution through the soil as silicic acid. Incomplete leaching allows for the precipitation of this silicic acid in the pedogenic zone. During drying induced by evapotranspiration, SiO4 is adsorbed onto surfaces, forming amorphous opaline silica. This process led to the silica cementation of south-facing terrace edges on the central California coast.[7] Alternatively, in situ alteration of eolian dust to opal-A, rather than precipitation from silica in soil solution, has been proposed for duripans in Idaho.[8] The eolian dust comprised volcanic glass and feldspars. Conditions favoring the precipitation of silica from solution include pH values less than 7, a high available surface area within the soil fabric, and high ionic strength of the soil solution. Although a high available surface area favors precipitation, silica cementation is most common in medium-textured parent materials with abundant skeletal grains. The grain-to-grain cementation is an important component in the formation of a duripan.[6] Aluminum- and Fe-oxyhydroxides specifically adsorb soluble silica. Since these compounds are nearly ubiquitous in soils, often as coatings on other grains, they form a significant sink for the precipitation of silica on surfaces.[3]

SILICA MORPHOLOGY Stages of silica cementation have been defined for coarse grained deposits in arid climatic regimes.[9] The first stage is precipitation of silica on the undersides of clasts. These are termed pendants, and are equivalent to opal beards in Australian soil descriptions.[10] Dramatic examples of opal pendants have been described in some central Californian soils.[11] The second stage is precipitation within the matrix, 1251

1252

where silica forms bridging contacts between grains. The third stage is silica precipitation on all sides of the clasts, and the final stage is a laminar cap above a plugged matrix. Apart from silica precipitation on gravels, silica can cement the finer matrix materials into nodules termed durinodes. Duripans are defined as soil horizons that are irreversibly cemented by various forms of silica such that they do not slake after soaking in water or hydrochloric acid.[12] Duripans take two general forms. In Mediterranean climates, duripans often have prismatic structure and pedogenic calcite is minimal or absent.[1] Small opal flocs within the matrix of the prisms are the

Pedogenic Silica Accumulation

precursors to durinodes. Further development of the duripan yields extensive grain-to-grain cementation in the soil matrix (Fig. 1) and silica coatings on the tops and upper sides of the prisms. Ultimately, silica laminae with platy structure cap the prismatic horizon.[1] This is in contrast to duripans that form in arid environments, which typically have a large component of pedogenic calcite and are characterized by a platy structure throughout, with plates 1–15 cm thick.[13] As little as 10% Si as opaline silica is adequate to cement horizons effectively, and form a fully developed duripan.[1] Thorough cementation of a soil horizon to form a duripan requires geomorphic stability of long duration, particularly in the absence of readily soluble volcanic ash. The ancient landscapes of Australia fit this criterion. Silica cemented horizons in Australia are common and are referred to as silcrete, duricrust, and grey billy.[10] Silica cementation has also been reported in hardsetting soils. Hardsetting soils have one or more horizons with hard to very hard consistence when dry. This phenomenon is recognized in soils with alternating seasons of wetting and drying, primarily those in Australia. Hardsetting soils eventually slake on wetting, and are thus not irreversibly cemented.[14] Amorphous particles and coatings of silica have been documented in these soils,[14] suggesting that silica is an important constituent in the process of seasonal cementation.[15] The easily mobilized silica might be ephemeral, and may or may not be related to incipient formation of a duripan.

REFERENCES

Fig. 1 Scanning electron micrograph of soil fabric in a duripan from southern California. (A) Primary mineral grains (P: plagioclase; B: biotite; Q:, quartz) cemented by opaline silica (O: opaline silica). (B) Close-up view of the lower left portion of Fig. 1A, showing biotite grains embedded in opaline silica. (Photo by Dr Krassimir Bozhilov, Central Facility for Advanced Microscopy and Microanalysis, UC Riverside.)


1. Flach, K.W.; Nettleton, W.D.; Gile, L.H.; Cady, J.G. Pedocementation: induration by silica, carbonates, and sesquioxides in the quaternary. Soil Sci. 1969, 107, 442–453. 2. Chadwick, O.A.; Hendricks, D.M.; Nettleton, W.D. Silica in duric soils: 1. A depositional model. Soil Sci. Soc. Am. J. 1987, 51, 975–982. 3. Dress, L.R.; Wilding, L.P.; Smeck, N.E.; Senkayi, A.L. Silica to soils: quartz and discordered silica polymosphs. In Minerals in Soil Environments, 2nd Ed.; Dixon, J.B., Weed, S.B., Eds.; Soil Sci. Soc.: Am. Madison, WI, 1989; 913–974. 4. Cleary, W.J.; Conolly, J.R. Embayed quartz grains in soils and their significance. J. Sediment. Petrol. 1972, 42, 899–904. 5. Wilding, L.P.; Drees, L.R. Contributions of forest opal and associated crystalline phases to fine silt and clay fractions of soils. Clays Clay Min. 1974, 22, 295–306. 6. Moody, L.E.; Graham, R.C. Silica-cemented terrace edges, central California coast. Soil Sci. Soc. Am. J. 1997, 61, 1723–1729.

Pedogenic Silica Accumulation

7. Blank, R.R.; Fosberg, M.A. Duripans in Idaho, U.S.A. in situ alteration of eolian dust (loess) to an Opal-A=XRay amorphous phase. Geoderma 1991, 48, 131–149. 8. Norton, L.D. Micromorphology of silica cementation in soils. In Soil Micromorphology: Studies in Management and Genesis; A.J. Ringrose-Voase and G.S. Humphreys, Eds.; Proc. IX Int. Working Meeting on Soil Micromorphology, Townsville, Australia, July 1992; Dev. Soil Sci., 1994; Vol. 22, 811–824. 9. Harden, J.W.; Taylor, E.M.; Reheis, M.C.; McFadden, L.D. Calcic, gypsic and siliceous soil chronosequences in arid and semiarid environments. In Occurrence, Characteristics, and Genesis of Carbonate, Gypsum, and Silica Accumulations in Soils; Nettleton, W.D., Ed.; Soil Sci. Soc. of Am. Spec. Publ.: Madison, WI, 1991; Vol. 26, 1–16. 10. Milnes, A.R.; Wright, M.J.; Thiry, M. Silica accumulations in saprolites and soils in South Australia. In Occurrence, Characteristics, and Genesis of Carbonate,


1253

11.

12. 13.

14.

15.

Gypsum, and Silica Accumulations in Soils; Nettleton, W.D., Ed.; Soil Sci. Soc. of Am. Spec. Publ.: Madison, WI, 1991; Vol. 26, 121–149. Munk, L.P.; Southard, R.J. Pedogenic implications of opaline pendants in some California late-pleistocene palexeralfs. Soil Sci. Soc. Am. J. 1993, 57, 149–154. Soil Survey Staff. In Soil Taxonomy; U.S. Gov. Printing Office: Washington, DC, 1975. Chadwick, O.A.; Graham, R.C. Pedogenic processes. In Handbook Soil Science; Summer, M.E., Ed.; CRC Press: Boca Raton, FL, 2000; E-41–E-75. Chartres, C.J.; Norton, L.D. Micromorphological and chemical properties of Australian soils with hardsetting and duric horizons. Dev. Soil Sci. 1994, 22, 825–834. Chartres, C.J.; Kirby, J.M.; Raupach, M. Poorly ordered silica and aluminasilicates as temporary cementing agents in hard-setting soils. Soil Sci. Soc. Am. J. 1990, 54, 1060–1067.

Pedological Modeling Ronald G. Amundson University of California, Berkeley, California, U.S.A.

INTRODUCTION ‘‘The fascinating impressiveness of rigorous mathematical analysis, with its atmosphere of precision and elegance, should not blind us to the defects of the premise that condition the whole process’’ by T. C. Chamberlin,[1] commenting on Lord Kelvin’s (ultimately incorrect) calculation of the age of the Earth.

In pedology, and in other sciences, ‘‘models’’ are increasingly used as tools for understanding natural phenomena. But what is a model, how are models used in pedology, and how are models developed and modified?

OVERVIEW To begin, the term ‘‘model’’ has been used interchangeably in pedology with other concepts, sometimes leading to confusion or miscommunication. Very simply, a ‘‘model’’ has been described by some as ‘‘a form of highly complex scientific hypothesis,’’[2] that is ‘‘a simplified and idealized description or conception of a particular system, situation, or process (often in mathematical terms) that is put forward as a basis for calculations, predictions, or further investigation.’’[3] As briefly outlined below, the mathematical approaches to describe a phenomenon can be varied, but all must rest on a solid understanding of the soil, and the factors and processes, which affect it. This empirical knowledge in turn constrains the ‘‘assumptions,’’ which underlie any mathematical model development. A model based on an incorrect, or poorly developed, understanding of a soil and the processes that affect it will likely be unable to describe the processes of interest, but that inability may in turn inspire the modeler to better understand the soil. Therefore, modeling can help refocus attention to fieldwork and to the type of data to be collected. A specific example of how assumptions affect the development of a model is given later in this entry. The first step in modeling soils—or anything—is to define the object of interest. In applying models to pedology, it should be recognized that soil is, in reality, a continuum of objects distributed across the earth’s surface—both in space and time. The exact lateral 1254 Copyright © 2006 by Taylor & Francis

boundary between one ‘‘soil’’ and another, or the vertical boundary between soil and nonsoil, is arguably impossible to determine. Jenny[4] first applied principles derived from the physical sciences to the conceptualization and modeling of soils. Jenny’s approach was to divide the continuum of soils on the earth’s surface into ‘‘systems,’’ which are arbitrarily defined, discrete, three-dimensional segments of the landscape that are amenable to mass or energy budgeting. The volume of these systems (both the chosen area and depth) is arbitrary, but it sets the stage for the mathematical formulations that are chosen to represent or describe it. A second important aspect of soils is the vast amount of time, and the array of unknown processes, that may have affected any soil system. This complex history in turn forces pedologists to develop tools, concepts, and modes of enquiry not always confronted by their experimental colleagues. Finally, soil formation is the result of an incompletely understood array of processes, and pedological models invariably are attempts to mathematically capture one, or at the most, a very restricted subset of the full suite of biogeochemical processes that affect a soil system. In the pedological literature, more attention has possibly been devoted to classifying and discussing models[5] than actually developing them. Unfortunately, much discussion has focused on the ‘‘pros and cons’’ of the Factors of Soil Formation. A System of Quantitative Pedology[4] in the realm of pedological models. Briefly, the factorial ‘‘model’’ discussed in a book-length treatise by Jenny[4] can be symbolically represented as s ¼ fðcl; o; r; p; t; . . .Þ where s is the soil properties; cl, climate; o, biota; r, topography; p, parent material; and t, time.[4,6] The general truthfulness of this statement is almost beyond dispute (virtually every pedologist would ultimately have to agree that soil forms in response to variations in these factors, and that soil properties can also be numerically correlated with variations in these factors). Indeed, soil is ‘‘defined’’ in terms of this statement: Soil is the ‘‘collection of natural bodies occupying portions of the earth’s surface that support plants and that have properties due to the integrated effect Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042725 Copyright # 2006 by Taylor & Francis. All rights reserved.

Pedological Modeling

1255

of climate and living matter, acting upon parent material, as conditioned by relief, over periods of time.’’[7]

Based on this definition, it would be more correct to define the state factor ‘‘model,’’ or ‘‘theory’’ as it has been alternately called, as a pedological ‘‘law’’ given its universal truthfulness and common definitions of scientific laws.[8] At the very least, it is a fundamental underlying theory of pedology in the sense of Kuhn.[9] This definition would also allow the discussion of models in pedology to move beyond this fundamental truism (and a useful quantitative and mathematical tool in its own right) to the development of mathematical models as they are commonly considered and used in geochemistry, geophysics, and related fields. The remainder of this entry deals with these more practical modeling issues.

extended for use in stable isotopic studies of soil organic matter[11] and in modeling turnover times of soil organic carbon.[12] However, the analytical model, as developed by Jenny, assumes constant inputs with time (a restriction that he noted does not occur in all situations) and constant decomposition rates with time. It also assumes that all organic matter is homogeneous (and by implication, has the same decomposition rates). Work over the past 15 years in particular has revealed that soil organic matter (and even litter layers) can best be viewed as multiple pools of soil organic matter, each with their own characteristic input and decomposition rates. In simplest terms, multiple pool models of soil organic matter can be expressed as n X dC ¼ ðIi ki Ci Þ dt i¼1

ð3Þ

MODELING PROCESSES Typical approaches to mathematically modeling soil processes involve the development of a mass or energy balance model. The mathematics used will ultimately hinge upon one’s understanding of the soil properties, the processes that presumably control them, and how these processes may vary over time. Some of the simplest models may be analytical models with time independent variables. Alternative modeling approaches may involve the abandonment of time invariant parameters and ultimately, the incorporation of relatively random changes in the rate of the process and factors that affect it. As an example, I begin with possibly the first true mathematical model of a pedogenic process—a time dependent, analytical, mass-balance model of O horizon formation in forest soils developed by Jenny, Gessel, and Bingham.[10] Jenny, Gessel, and Bingham[10] defined the system of study, discussed the processes that affect it, and, for the simplest cases they considered (tropical forests with nearly constant litter inputs with time), described the change in O horizon mass (F, in mass per area) with time dF ¼ Adt kðF þ AÞdt

ð1Þ

where A is litter inputs (mass=area=time), and k, decomposition constant (1=time) (Ref. [10] gives a fuller discussion of calculation and definition of k). Upon integration, the solution to Eq. (1) provided by Jenny was F ¼

Að1 kÞ ð1 ekt Þ k

ð2Þ

This model, and permutations of it, has served as the foundation for decades of research and modeling of the soil organic C budget. Today the model has been


where C is the total mass of soil C (mass=volume); I, inputs of pool i (mass=volume=time); ki, decomposition constant of pool i (1=time); and Ci, mass of soil C in pool i (mass=volume).[11] Even these multiple pool models do not capture other aspects of soil formation, the variation in soil C with depth, and the likelihood that the process may have varied unpredictably over time. To address the first issue, models may include a downward transport term and depth dependent inputs. For a single pool of organic matter, a basic depth dependent model is dC=dt ¼ I u

@C kC @Z

ð4Þ

where u is the advection coefficient (distance=time).[11] Even these models can be made more complex, for example to include the process of diffusion.[13] Analytical solutions to all the aforementioned models require time invariant parameters. The inclusion of time-dependent parameters may be accomplished through numerical means. Yet, soils form over thousands to millions of years, with numerous unknown perturbations to the system that elude the fundamental framework of these simple models.[14] It is these real world complications imposed by nature that weaken the utility of relatively simple, time invariant, deterministic models. Recently, Phillips[14] began the discussion of applying nonlinear dynamical system theory to soil modeling, with the goal of explicitly incorporating the role that random differences in initial conditions and historical contingencies have on the processes of soil formation. Undoubtedly, work of this nature is one of the future challenges in the field of pedological modeling.

1256

CONCLUSIONS While models are powerful means of testing hypotheses and synthesizing contemporary knowledge in a concise way, the previous discussion serves to illustrate that models of all types are simplified, incomplete, mathematical descriptions of the ‘‘real’’ world. Increasingly, it is emphasized that models can never be fully verified (confirmed as the establishment of truth).[2] Experience shows that multiple models (or versions of the same model) may faithfully mimic empirical observations of interest, a dilemma illustrating a practical verification problem. In addition, mathematical models may be internally correct but they may poorly represent the phenomena they intend to describe because of incomplete knowledge of the system and, as a result, incorrect assumptions. Knowledge of the soil is essential in modeling, for as Baker[15] has recently noted, ‘‘mathematics (is) the science that draws necessary conclusions without regard to facts.’’ Given the ultimate simplicity and abstractness of models, and the ultimate complexity of nature in general, and soils in particular, the question ‘‘why model in the first place?’’ may be asked.[2] Pedological processes are first order controls on global atmospheric[16] and aquatic chemistry, and even relatively simple analytical models of soil processes have thus far proven useful to link these global reservoirs. The truly unique role for pedologists, in addition to applying mathematics on their own, is to provide the unique conceptual foundation peculiar to soils—one informed by extensive field observations guided by ideas generated during previous modeling attempts—that will make pedological models relevant to scientists and society.

REFERENCES 1. Chamberlin, T.C. On Lord Kelvin’s Address on the Age of the Earth as an Abode Fitted for Life; Smithsonian Institution Annual Report: Washington, DC, 1899; 223–246.


Pedological Modeling

2. Oreskes, N.; Shrader-Frechette, K.; Belitz, K. Verification, validation, and confirmation of numerical models in the earth sciences. Science 1994, 263, 641–646. 3. The Oxford English Dictionary, 2nd Ed.; Clarendon Press: Oxford, 1989. 4. Jenny, H. Factors of Soil Formation. A System of Quantitative Pedology; McGraw Hill Book Co.: New York, 1941; 281 pp. 5. Hoosbeek, M.R.; Amundson, R.; Bryant, R.B. Pedological modeling. In Handbook of Soil Science; Sumner, M.E., Ed.; CRC Press: Boca Raton, FL, 1999; E77–E116. 6. Amundson, R.; Jenny, H. On a state factor model of ecosystems. Bioscience 1997, 47, 536–543. 7. Soil survey staff. Soil Survey Manual; U.S. Dep. Agric. Handbook No. 18; U.S. Govt. Printing Office: Washington, DC, 1951; 503 pp. 8. Morris, C., Ed. Academic Press Dictionary of Science and Technology; Academic Press: New York, 1992. 9. Kuhn, T.S. The Structure of Scientific Revolutions, 2nd Ed.; University of Chicago Press: Chicago, 1970; 210 pp. 10. Jenny, H.; Gessel, S.P.; Bingham, F.T. Comparative decomposition rates of organic matter in temperate and tropical regions. Soil Sci. 1949, 68, 419–432. 11. Amundson, R.; Baisden, W.T. Stable isotope tracers and models in soil organic matter studies. In Methods in Ecosystem Science; Sala, O., Mooney, H., Howarth, B., Jackson, R.B., Eds.; Springer Verlag: New York, 2000; 117–137. 12. Trumbore, S.E. Comparison of soil carbon dynamics in tropical and temperate soil using radiocarbon measurements. Global Biogeochem. Cycles 1993, 9, 515–528. 13. Elzein, A.; Balesdent, J. Mechanistic simulation of vertical distribution of carbon concentrations and residence times in soils. Soil Sci. Soc. Am. J. 1995, 59, 1328–1335. 14. Phillips, J.D. On the relations between complex systems and the factorial model of soil formation (with discussion). Geoderma 1998, 86, 1–42. 15. Baker, V.R. Geosemiosis. Geol. Soc. Am. Bul. 1999, 111, 633–645. 16. Amundson, R.; Stern, L.; Baisden, T.; Wang, Y. The isotopic composition of soil and soil-respired CO2. Geoderma 1998, 82, 83–114.

Pedotransfer Functions J. H. M. Wo¨sten Alterra Green World Research, Wageningen, The Netherlands

INTRODUCTION Simulation models, which are indispensable tools in modeling water and solute movement into and through soil, require as key input parameters easily accessible and representative hydraulic characteristics. Techniques to measure these characteristics are relatively time-consuming and therefore costly. At the same time, good predictions of the characteristics instead of direct measurements may be accurate enough for many applications. Considering the desired accuracy and the available financial resources, it is rewarding to analyze existing databases containing measured hydraulic characteristics and to establish relationships that predict the characteristics from measured basic soil data. These predictive relationships are called ‘‘pedotransfer functions’’ (PTFs)[1] and they essentially translate data ‘‘we have’’ into data ‘‘we need.’’ Basically, PTFs relate soil characteristics being assembled during soil survey to more complex characteristics needed for simulation. Predicting soil hydraulic characteristics dominates the research field, though soil chemical and soil biological characteristics are also being predicted. Several reviews on PTF development and use have been published.[2,3] Large databases on measured hydraulic characteristics, such as UNSODA,[4] HYPRES,[5] WISE,[6] and United States Department of Agriculture Natural Resource Conservation Service pedon database,[7] form the essential, basic sources of information for the derivation of PTFs. In using PTFs, insight is needed to determine the input variables that are to be included in a PTF, what technique is to be used to establish a PTF, and how accuracy and reliability of PTFs are to be quantified.

FUNCTIONS USED TO DESCRIBE THE WATER RETENTION AND HYDRAULIC CONDUCTIVITY CHARACTERISTICS Describing hydraulic characteristics as functions rather than as tables has the clear advantage that they can be easily incorporated in simulation models. There exists a wide range of different equations for the description of the characteristics. The following equations to describe volumetric soil water content, y, and hydraulic Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042726 Copyright # 2006 by Taylor & Francis. All rights reserved. Copyright © 2006 by Taylor & Francis

conductivity, K, as functions of pressure head, h, are widely used[8] yðhÞ ¼ yr þ

KðhÞ ¼ Ks

ys yr ðl þ jahjn Þ11=n

½ðl þ jahjn Þ11=n jahjn1 2 ðl þ jahjn Þð11=nÞðlþ2Þ

ð1Þ

ð2Þ

In these equations, the subscripts r and s refer to residual and saturated values and y, n, and l are parameters that determine the shape of the curve. The residual water content yr refers to the water content, where the gradient dy=dh becomes zero (h ! 1). The parameter a (1=cm) approximately equals the inverse of the pressure head at the inflection point. The dimensionless parameter n reflects the steepness of the curve. The dimensionless parameter l determines the slope of the hydraulic conductivity curve in the range of more negative values of h. Pedotransfer functions to predict the model parameters yr, ys, Ks, a, l, and n from basic soil data were built by many authors.[9] Fig. 1 shows the mean water retention and hydraulic conductivity characteristics, also called class PTFs, for the texture class ‘‘medium fine topsoil.’’[5] SELECTION OF PEDOTRANSFER FUNCTION PREDICTOR VARIABLES Soil properties affecting water retention and hydraulic conductivity are manifold.[10] Table 1 lists the properties used most often as predictors because of their availability and because they proved to be the most promising ones. ‘‘Particle size distribution’’ is used in almost all PTFs. Particle size classes differ in different national and international classification systems, and so the number and the size of classes used in PTFs may also differ. Using sand, silt, and clay contents is a common approach. ‘‘Limited, measured water retention data’’ at, for instance, two pressure heads may dramatically improve predictions of the complete water retention characteristic. ‘‘Porosity or bulk density’’ is an important variable in many PTFs. 1257

1258

Pedotransfer Functions

Fig. 1 Geometric mean water retention (A) and hydraulic conductivity (B) characteristic solid lines, standard deviations bars, and van Genuchten fits dotted lines for the texture class ‘‘medium fine topsoil.’’

‘‘Soil structure and morphology descriptors’’ such as the parameter topsoil and subsoil are successfully included in PTFs. ‘‘Landscape position’’ is used as a topographic variable in PTFs. ‘‘Organic matter content’’ is often used as predictor, but because bulk density and organic matter content are correlated, bulk density may effectively substitute organic matter content. ‘‘Mechanical properties and shrink–swell parameters,’’ as characterized by the coefficient of linear

extensibility (COLE), are used to estimate both water retention and Ks.

METHODS TO DEVELOP PEDOTRANSFER FUNCTIONS When the set of PTF input parameters is defined and the PTF output is decided upon, a method is selected to relate input and output. The most prominent methods to create this relationship are discussed below.

Table 1 Soil properties often used in PTFs Particle size properties Sand, silt, clay

Hydraulic characteristics

Chemical/mineralogical properties

Bulk density

Organic carbon

33 kPa

Porosity

Organic matter

1500 kPa

Horizon

CEC

Structure

Clay type

Median or geometric mean particle size

Grade Size Shape

CaCO3 Iron

Water-stable aggregates

Color Consistence Pedality Landscape position

Fine sand Very coarse sand, coarse fragments


Water content at

Morphological properties

Reference moisture retention curve

Mechanical properties Penetration resistance

Pedotransfer Functions

Table 2 Continuous PTFs developed from the HYPRES database ys ¼ 0.7919 þ 0.001691C 0.29619D 0.000001491S2a þ 0.0000821OM2 þ 0.02427C1 þ 0.01113=S þ 0.01472ln(S) 0.0000733OMC 0.000619DC 0.001183DOM 0.0001664 topsoil S (R2 ¼ 76%) a ys is the saturated water content in the van Genuchten equations; C, percentage clay (i.e., percentage 8.7. A very high pH indicates the presence of magnesium carbonate.[9] Encyclopedia of Soil Science DOI: 10.1081/E-ESS-120042727 Copyright # 2006 by Taylor & Francis. All rights reserved.

Petrocalcic Horizons, Soils with

Fig. 1 Below an A horizon (1) is a layered noncompact petrocalcic horizon (crust) (2), over an unlayered (nonplaty) petrocalcic horizon (3), on a calcic horizon (4).

Soils with petrocalcic horizons in environments with restricted drainage tend to form montmorillonite, attapulgite, and sepiolite clay minerals.[6–8,10]

VERTICAL SUCCESSION OF THE CALCIC AND PETROCALCIC HORIZONS Below A or B horizons, the accumulation of calcium carbonate can be of three types: Slightly differentiated: the distribution of carbonates is diffuse, sometimes with pseudomycelia with very diffuse upper and lower limits. Moderately differentiated: the carbonates in the calcic horizon occur partly as diffuse distributions in the soil matrix and partly as concentrations, forming cutans, soft and hard nodules, or veins, with diffuse upper and lower limits. Highly differentiated (petrocalcic): the carbonates form laterally continuous concentrations


1267

Fig. 2 A layered compact petrocalcic horizon (slab) (2) below an A horizon (containing pebbles of slab) (1) over a layered noncompact petrocalcic horizon (crust) (3) on a calcic horizon (4).

resulting in one or more superposed petrocalcic horizons. The vertical complete succession from the top is: finely layered, over thicker compact layers (slabs), over thick noncompact layers, over a massive horizon. There are four main types of vertical successions of petrocalcic horizons: – single massive horizons; – layered noncompact horizons, over unlayered horizons; – finely layered horizons, over layered noncompact horizons, over unlayered horizons; – finely layered horizons, over layered compact horizons, over layered noncompact horizons, over unlayered horizons. Petrocalcic horizons always have sharp upper boundaries and carbonate contents that decrease with depth. A gradual transition and discontinuous

1268

concentrations of carbonate mark the lower boundary of the petrocalcic horizon. The solum above the petrocalcic horizon is generally between 10 and 50 cm thick.

LATERAL DISTRIBUTION OF THE CALCIC AND PETROCALCIC HORIZONS As with vertical transitions, progressive lateral transitions frequently occur between different forms of calcic and petrocalcic horizons. From this, it can be concluded that the vertical and horizontal structures of calcium carbonate accumulation result from the same mechanisms.


In a landscape, the lateral distribution of the calcic and petrocalcic horizons is mainly a function of topography and age. Along a slope or pediment, calcic horizon development increases downslope. In a complete toposequence (catena), differentiation progresses from a soil with minimal calcic horizon development upslope, to a soil with moderate calcic horizon development immediately downslope, and finally to soils with more and more strongly developed petrocalcic horizons, the different types and superposition of petrocalcic horizons appearing successively (Fig. 3). In time, calcic horizons evolve (Fig. 3) from diffuse carbonate distributions to discontinuous accumulations

Fig. 3 Relationships between calcic and petrocalcic horizons in space and time (North Morocco). The length of the sequences may vary between some tens and several hundreds of meters; the difference in altitude, between the old and recent Quaternary surfaces, is some tens of meters. (From Ref.[1].)



(pseudomycelia, cutans, nodules, or veins) to unlayered petrocalcic horizons to layered noncompact petrocalcic horizons to slab. The finely layered (ribboned) horizons can exist as soon as noncompact layered horizons appear. On the other hand, decarbonation of the A and B horizons above the calcic and petrocalcic horizons does not increase with age. It is only on the younger recent Quaternary surfaces that a small decarbonation can be observed as these surfaces age. However, this decarbonation does not increase on the older surfaces and, when petrocalcic horizons appear, the calcium carbonate content of the upper horizons may increase due to erosion and subsequent formation of the upper horizons from the calcrete. These facts confirm the following interpretations: There is a logical order of appearance of the calcic and petrocalcic accumulations and horizons; this logical order is the same in space, vertically and laterally, and in time. These accumulations and horizons are thus genetically linked, by toposequences and chronosequences. The vertical leaching of the calcium carbonate, which impoverishes the upper horizons in favor of the calcic and petrocalcic horizons, is a limited phenomenon; the major part of the calcium carbonate that accumulates in the soils comes from lateral redistributions. In arid and semiarid regions, which are the privileged domains of the petrocalcic horizons, calcic and petrocalcic horizons can occur in soils formed from noncalcareous or noncalcic rocks: this happens when landscapes upstream furnish calcium by lateral lixiviation. However, in very arid regions near the sea, very strong petrocalcic horizons occur in soils formed from noncalcareous, noncalcic rocks without possible upstream sources of calcium. So calcium carbonate can also arrive by air as calcareous dust and calcium from sea spray.

SOIL USE Petrocalcic horizons in soils at depths exchangeable > organic matter bonded >

Table 2 Concentrations and coefficient of variation of rare earth elements in different sludges Night soil sludgea (n = 10) Elements

Mean (mg kg1)

CV (%)

Sewage sludge (n = 14) Mean (mg kg1)

CV (%)

Food industry sludge (n = 10) Mean (mg kg1)

Chemical industry sludge (n = 10)

CV (%)

Mean (mg kg1)

CV (%)

La

3.39

37

6.70

47

0.89

72

2.46

98

Ce

6.98

44

14.10

58

1.83

77

2.69

105

Pr

0.82

38

1.48

46

0.22

82

0.48

95

Nd

3.18

34

6.00

47

0.91

82

2.04

98

Sm

0.53

36

1.02

40

0.17

81

0.36

95

Gd

0.53

34

1.18

45

0.17

79

0.48

101

Tb

0.07

45

0.16

36

0.03

81

0.06

103

Dy

0.39

53

0.93

33

0.14

76

0.39

110

Ho

0.07

54

0.19

32

0.03

79

0.09

119

Er

0.21

55

0.57

31

0.08

73

0.26

118

Tm

0.03

52

0.08

26

0.01

81

0.03

112

Yb

0.20

58

0.54

31

0.09

83

0.19

109

Lu

0.03

56

0.08

31

0.01

84

0.03

105

a

Feces and urine of humans. (From Ref.[5].)


Rare Earth Elements

1439

Table 3 Mean content of REEs in soils extracted by Na2O2=NaOH (n ¼ 467)

Light REEs

Table 5 Total concentration of rare earth elements in plants (mg kg1)

Total contents (mg kg1) and ratios

Species

Content (mg kg1)

Heavy REEs

Content (mg kg1)

La

41.2

Gd

4.8

Total REE (T)

172.8

Ce

73.4

Tb

0.7

Light REE (L)

156.0

Pr

7.3

Dy

4.4

Heavy REE (H)

16.8

Nd

27.5

Ho

0.9

L=H

9.3

Sm

5.6

Er

2.7

L=T

0.9

Eu

1.1

Tm

0.4

Yb

2.5

Lu

0.4

(From Ref.[1].)

Fe=Mn oxide bonded REEs.[9] The formation of bridged hydroxo complexes is probably the dominant sorption mechanism to clay minerals.[14] Clay type, pH, CEC, organic matter, and amorphous iron content regulate the adsorption kinetics of REEs.[1,2,9] Langmuir and Freundlich equations were found to describe precisely the absorption of REEs in soils.[1,15]

Rice Wheat Corn Cucumber Leek Spinach Cauliflower Lotus root Tomato Chinese cabbage Pepper Potato Cabbage Mushroom Orange Litchi Grape Longan Banana Apple Pear Watermelon Sugarcane Peach

n

Min (mg kg1)

Max (mg kg1)

319 440 139 41 33 41 61 31 64 67 31 34 38 33 41 30 61 30 33 62 34 37 27 4

Encyclopedia of soil science, Volume 1

Encyclopedia of Radio, Volume 1

Fundamentals of Soil Science

Encyclopedia of the History of Arabic Science Volume 1

Cowboy Encyclopedia Volume 1

Encyclopedia of the History of Arabic Science Volume 1

Biology of soil science

Fundamentals of Soil Science

Environmental encyclopedia, Volume 1

The Encyclopedia of Political Science, Volume 1-5

Encyclopedia of Grain Science, Three-Volume Set (Vol 1-3)

Encyclopedia of Marine Science (Science Encyclopedia)

The Internet Encyclopedia, Volume-1

The Catholic Encyclopedia Volume 1

The Gale Encyclopedia of Cancer, volume 1

Encyclopedia of World History Volume 1

Encyclopedia of social measurement, Volume 1

Encyclopedia of Nanoscience and Nanotechnology Volume 1

Encyclopedia Of Mathematics (Science Encyclopedia)

Encyclopedia of Computational Mechanics: Volume 1 Fundamentals

An Encyclopedia of Freemasonry Volume 1

Encyclopedia Of Chemistry (Science Encyclopedia)

Encyclopedia Of Mathematics (Science Encyclopedia)

Encyclopedia Of Chemistry (Science Encyclopedia)

Encyclopedia Of Chemistry (Science Encyclopedia)

Soil Science Simplified

Fractals in Soil Science

Soil Heavy Metals (Soil Biology, Volume 19)

Soil Enzymology (Soil Biology, Volume 22)

Proximal Soil Sensing (Progress in Soil Science)

Encyclopedia of soil science, Volume 1

Encyclopedia of Radio, Volume 1

Fundamentals of Soil Science

Encyclopedia of the History of Arabic Science Volume 1

Cowboy Encyclopedia Volume 1

Encyclopedia of the History of Arabic Science Volume 1

Biology of soil science

Fundamentals of Soil Science

Environmental encyclopedia, Volume 1

The Encyclopedia of Political Science, Volume 1-5

Encyclopedia of Grain Science, Three-Volume Set (Vol 1-3)

Encyclopedia of Marine Science (Science Encyclopedia)

The Internet Encyclopedia, Volume-1

The Catholic Encyclopedia Volume 1

The Gale Encyclopedia of Cancer, volume 1

Encyclopedia of World History Volume 1

Encyclopedia of social measurement, Volume 1

Encyclopedia of Nanoscience and Nanotechnology Volume 1

Encyclopedia Of Mathematics (Science Encyclopedia)

Encyclopedia of Computational Mechanics: Volume 1 Fundamentals

An Encyclopedia of Freemasonry Volume 1

Encyclopedia Of Chemistry (Science Encyclopedia)

Encyclopedia Of Mathematics (Science Encyclopedia)

Encyclopedia Of Chemistry (Science Encyclopedia)

Encyclopedia Of Chemistry (Science Encyclopedia)

Soil Science Simplified

Fractals in Soil Science

Soil Heavy Metals (Soil Biology, Volume 19)

Soil Enzymology (Soil Biology, Volume 22)

Proximal Soil Sensing (Progress in Soil Science)

Recommend Documents