1. Soils, What They Are and How They Form

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1.1. Introduction

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Arizona, the sixth largest state, has an area of 295,146 square kilometers (113,956 square miles) or 29,515,111 hectares (72,931,840 acres). The state soil resources are the outermost surface of the land and range in depth from several centimeters (a few inches) to 1 or 2 meters (3 to 7 feet). The soils are among the most important state natural resources. Soils provide sustenance for animals and man and support buildings and highways, while contributing to the economies of our cities, the productivity of our farms and rangelands, and the vitality of our wildlife and wilderness areas.

Put differently by Jacks and Whyte ( 1939), ‘‘Below that thin layer comprising the delicate organism known as the soil is a planet as lifeless as the moon.’’ Either statement acknowledges an awareness of the necessity of soil in our lives. It was important too between 12,000 and 15,000 years ago to the first humans who entered what is now called Arizona. They must have been aware of the support soil gave to the various kinds of life forms that they encountered in their travels, of the interactions between soil and the life forms.

To these people the most important life forms were faunal and it probably was pursuit of animals that brought them to this part of the world, perhaps from Asia. Earliest Arizonans moved with the animals that sustained them, leaving an occasional glimpse of their presence here in prehistoric campsites and the remains of elephants, camels, horses, sloths and bison that they killed.

As the climate changed and as the big game became less plentiful, early Arizonans sustained themselves by hunting the smaller species and gathering food from plants. These hunter-gatherers lived here for perhaps 8,000 years and are known to researchers as the Cochise people because the first evidence of them was found in Cochise County.

The region experienced climatic shifts over time, however, to produce a rich ecological system for sustaining ancient animals, and early man. Many plants, like paloverde and mesquite, produced abundant protein sources for food consumption as well as nitrogen that aided general plant growth in the system. This environmental richness before the time of Christ was extended by the adoption of farming practices.

Evidence of the first farmers in the region has been found in the rich alluvial valleys of Cochise County dating from around 2,000 BC. At this time, corn or maize was first domesticated as a reliable food source. Ancient corn had separately sheathed kernels in a husk, attached to a small cob. Acceptance and cultivation of this ancestor of corn transformed human social and economic life from hunting and gathering to farming and trading. The simultaneous introduction of beans, chili and squash, along with corn, provided complete nutrition for early man. With this storable source of food available in sufficient, and later surplus, quantities, the small bands of humans became larger in numbers, more elaborate in organization and more concentrated in definite centers. Later, trading food for specialized finished goods, like fabrics and pottery, led to development of sophisticated societies.

Archaeologists have noted that distinctions in geographic regions began to emerge over time. The Hohokam, Mogollon and Anasazi became unique cultural entities, with identifiable traits. The Hohokam were in the riverine valleys. The Mogollon were in areas of the mountains and the Mogollon

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Rim. The Anasazi were in the far north on the high plateau. All these cultures had economic and social distinctions and attained their highest levels of development in the period from 900 AD to 1450 AD, a wetter period. However, the Hohokam, Mogollon and Anasazi did share numerous traits.

All three cultures seem to have been stimulated by the introduction of new crops, agricultural practices and social customs that probably came from highly developed cultures in Mexico (University of Arizona Press, 1972). This stimulation apparently led to trade, as evidenced by discovery of pottery designs traditional to northern peoples in southern sites and by traditional legends of south-to-north movements.

The agricultural practices adopted by these people and their successors had to be refined in each region and adapted to the constraints of soils and water. Mountain and plateau people were essentially dryland farmers, depending on rainfall and soil moisture. Arable lands were limited by the hilly and mountainous terrain. Desert-dweller domains were limited by the sources of water, whether flowing rivers or runoff from rains. Desert dwellers practiced irrigation farming in river valleys and water harvesting to augment dryland farming.

The Hohokam built large irrigation canals that drew water from flowing streams. Several hundred miles of farm engineering systems can be seen today in the Salt and Gila river valleys. Some match modern systems in size. There is evidence that pioneers restored Hohokam irrigation ditches by patching and cleaning them and then used them to irrigate their crops. Certainly, no Indian achievements north of central Mexico in pre-Conquest times surpassed the Hohokam canal system. They stand as examples of sound planning, enormous expenditures of effort and evident intercommunity organization. The Hohokam were master farmers, producing corn, beans and squash, as well as a unique contribution, cotton.

FIGURE 1. Indian Reservations in Arizona

Outward from the Hohokam agriculture centers were other desert dwellers, living on areas now overlapping the current Papago Reservation. Here, people did not have access to live streams. Water-harvesting techniques were perfected to divert runoff from rainfall. One water-harvesting gathering or diversion ditch ran in a westerly direction for nearly 16 km (10 miles) from the base of Baboquivari Peak. It cut across numerous small natural drainage washes on the gentle piedmont slope, collecting rain and runoff, directing it to fields in fertile lowland ground (University of Arizona Press, 1972).

The Mogollon people lived in high areas that ranged from the Mogollon Rim and White Mountains southward into Mexico. Remains of their pueblo-style dwellings are found surprisingly close to Hohokam sites, often within close walking distances, but on high ground. While water may not have been a problem for the Mogollon, arable land was. Trade with the Hohokam can be seen in Hohokam sites. The Mogollon evidently specialized in pottery production, trading with the Hohokam, who could grow surpluses of food and fiber. Among the earliest people to leave evidence of cultivated corn, the Mogollon became a part of this regional mosaic of cultures (University of Arizona Press, 1972).

Farther north were the Anasazi, inhabiting the high plateau region, where running streams were scarce. Populations were concentrated into large centers at Mesa Verde, Chaco Canyon, Hopi and Black Mesa. In the White Mountains, the Anasazi and Mogollon evidently overlapped. The diversified environment and terrain of the Anasazi region demanded development of various methods for cultivating limited arable soils to produce food to support these large communities. Techniques included intensive cropping of carefully cultivated arable soils and water harvesting of rain and floodwater.

During the first Spanish explorations of the Southwest, between 1535 and 1604 AD, the descendants of the Hohokam, Mogollon and Anasazi offered their visitors food, not the gold Spanish tales predicted. Food included, corn, beans and squash, also found in archaeological sites dated from earlier times. Most of the tribes met by the Spanish are still represented in Arizona (See Figure 1).

The study of soils is now the occupation of the soil scientist. An important aspect of soil science is the study of the soil as a natural phenomenon on the surface of the Earth. The soil scientist, then, is interested in the appearance of the soil, its mode of formation, its physical, chemical and biological processes and composition, and its classification, distribution and use.

A modern soil scientist, referring to Plate 1, might tell us for instance that the Hohokam probably farmed in HA1, HA2,

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HA3, HA5 and HA8 as well as in some areas in TS2, TS3 and TS19; that the Mogollon probably farmed only in TS9 and TS12 in the northeastern area of the Thermic Semiarid Soils; and that the Anasazi probably farmed mostly in MA5 and MA6 and some in MA1, MA2, MA3 and MA4.

Soil science is an integrative science. It uses many branches of scientific knowledge, including certain aspects of physics, chemistry, biology and geology as well as agriculture, forestry, history, geography and archaeology. From all of these, information is obtained that can be synthesized into a scientific discipline and natural philosophy separate from, and yet closely related to, many other branches of natural science.

Soil science can be viewed as a pure science that investigates the basic processes involved in the formation of soils and that produces soil maps and classifications. The scientific knowledge obtained is extremely important for those who must solve practical problems in agriculture, horticulture, range management, forestry and plan the use of the land.

1.2. Definition of Soil

The term soil generally is used to describe the material on the thin skin of the Earth's crust and that has been under the influence of certain physical and biological processes. Soil is considered here primarily as a natural unit in a pedological or ecological sense rather than in an engineering sense.

Review of detailed soil maps from many different regions of the world reveals a pattern of the interplay of climate with other soil forming factors. Soils, then can be considered to be in a state of dynamic equilibrium or slow evolution with their environments.

Soil also can be envisaged as an open system through which various hydrological, biological and geochemical cycles operate, a process-response system in which exists a close relationship between soil properties and the inputs and outputs of mass and energy.

FIGURE 2. Generalized Soil Composition Diagram

In addition to these concepts, soil is considered in this book according to the definition given in Soil Taxonomy (Soil Survey Staff, 1975): ‘‘Soil is the collection of natural bodies on the earth's surface, in places modified or even made by man of earthy materials, containing living matter and supporting or capable of supporting plants out-of-doors. Its upper limit is air or shallow water. At its margins it grades to deep water or to barren areas of rock and ice.... Commonly soil grades at its lower margin to hard rock or to earthy materials virtually devoid of roots, animals or marks of other biological activity.’’

1.3. Composition of Soils

Soils have four main constituents: mineral and organic matter, air and water (Figure 2). Mineral matter includes all minerals inherited from the parent material as well as those formed by recombination from substances in the soil solution. Organic matter is derived mostly from decaying plant material broken down and decomposed by the actions of animals and microorganisms living in the soil. Normally, both air and water fill the voids in soil.

Minerals are the major constituent in Arizona soils and are derived from the parent material by weathering. Mineral particles range in size from 2.0 mm to less than 0.002 mm (0.079 to 0.000079 in). These particles constitute the fine earth of soil and are the bases upon which soil texture is determined according to the relative amounts of the various particles in the soil. Soil minerals are sand, particles that range in size from 0.05 to 2.0 mm (0.002 to 0.079 in) in diameter; silt, particles having diameters of 0.002 to 0.05 mm (0.000079 to 0.002 in); and clay, particles less than 0.002 mm (0.000079 in) in diameter (Figure 3).

FIGURE 3. Representative Shapes and Sizes of Sand, Silt and Clay

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FIGURE 4. Soil Texture Triangle

The textural class of a soil can be determined by using the soil textural triangle (Figure 4). A soil with a loam texture will have the proportions of clay, silt and sand to place it in the ‘‘loam area’’ of the triangle. With practice, the texture of a soil can be approximated in the field by moistening and then working a sample between the fingers and thumb. Based upon how the soil feels, the porportions of clay, silt and sand is approximated. Larger particles such as gravel and other rock fragments also occur in soil, but are considered inert, contributing only their bulk and physical presence to the soil.

Minerals can be inherited directly from the parent material with little or no change or they may be formed from minerals within the parent material during weathering and soil formation. Primary minerals such as quartz, feldspars, various mafic minerals, muscovite and biotite commonly are inherited from the parent material, whereas clay minerals form as weathering products in the soil. Generally, the proportion of primary minerals decreases and the proportion of clay minerals increases as the soil particle size becomes smaller Thus, the sand- and silt-size particles are dominated by primary minerals and the clay by clay minerals.

Clays are the most important mineral constituents of soils Because of their small size, clay minerals have large specific surface areas (surface area per unit mass) compared with the larger silt and sand particles. Most chemical and physical reactions and interactions occur on the surfaces of soil particles. Therefore, whether chemical or physical phenomena are considered, the small clay particles with their large specific surface area are most important in determining the fundamental properties of soil.

Clay minerals are characterized by a layered crystalline structure. The clay minerals mica and montmorillonite dominate the clay in most soils in Arizona. A crystal unit of mica, such as muscovite, has one silica sheet on each side of an alumina sheet. Adjacent crystal units are held together by potassium bridges so that the space between the units is not readily accessible for surface reactions (Figure 5). The crystal structure of montmorillonite is similar to that of mica but the adjacent crystal units are bound together by very weak linkages (Figure 6). The entire surface of the montmorillonite crystal, therefore, is accessible for reactions. Distances between the crystal units are determined by the amount of water present. As amounts of water vary, shrinking and swelling occurs with drying and wetting. Kaolinite is a common component in many of the soils in Arizona, although usually in relatively small amounts. The structure of kaolinite differs from that of muscovite and montmorillonite; it has only one silica sheet bonded to one alumina sheet (Figure 7). Kaolinite particles consist of these crystal units stacked on each other and are bonded so that the interlayer positions are not accessible for surface reactions.

FIGURE 5. Crystal Structure of Muscovite Clay (after P. W. Birkland, 1974)

A very important and intriguing property of clay minerals is their ability to adsorb and hold bases such as calcium, magnesium, sodium and potassium, and the acid element hydrogen. Clay minerals, because of their chemical composition and structure, possess a net negative charge. The bases, also known as cations, are positively charged when in the soil solution and are attracted to and held on the surfaces of the negatively charged clay particles. The adsorbed cations are held tightly enough to retard their movement from the soil by leaching, yet loosely enough to be replaced by other cations, a process called cation exchange. The absolute quantity of cations that may be held in a soil by the clay fraction in exchangeable form is the cation exchange capacity. One of the most significant features of the cation exchange capacity of a soil is that it provides temporary storage of large quantities of plant nutrients such as calcium, magnesium and potassium. When ammonium sulfate fertilizer is added to a soil it dissolves in the soil solution

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forming the ammonium cation and the negatively charged sulfate ion. The ammonium then may replace some of the exchangeable cations and be held on the clay, silt and sand surfaces until used by plants. This reaction occurs most readily on clay particles, to a lesser extent on silt and to a much lesser extent on sand.

FIGURE 6. Crystal Structure of Montmorillonite Clay (after P. W. Birkland, 1974)

Arizona soils generally contain relatively small amounts of organic matter due to limited plant growth and rapid decomposition of dead plant matter. Most of the soils contain less than 1 percent organic matter, although a few high-elevetion meadow soils in the White Mountains contain more than 10 percent. Organic matter present in soils is intimately mixed with the mineral matter of the surface horizons and it may lie on the soil surface. Surface organic matter accumulations are quite limited in the deserts of the state but are very common in the forested soils. This surface organic matter may be freshly deposited leaf and other plant debris, as well as matter in different degrees of decomposition. Organic matter also adsorbs and holds ions in exchangeable forms similar to the clay minerals.

Soil structure is an important physical characteristic of any soil. It is produced by the aggregation of particles of sand, silt and clay into larger units called peds. Formation of structure is enhanced by organic matter, clay, CaCO3 (calcite), iron oxides and other substances that can bind particles together. The five classes of soil structure are based on the geometrical shape of the peds (Figure 8): platy, granular, blocky, prismatic and columnar structures. Platy and granular aggregates usually are on the soil surface or in the surface horizon. Blocky, prismatic and columnar agregates usually occur in the B horizon, with sodium affecting the columnar structure. Soils that lack structures are called structureless. These include soils composed of loose sand grains and those in which the particles form a coherent mass, but do not break out into distinct peds.

Air and water fill voids of the soil. Voids may result from the way soil particles in the form of single grains or peds are packed together. They are called packing voids and normally are connected. Voids other than packing voids may exist that are not normally interconnected with other voids. These voids may have irregularly shaped walls or walls consisting of smooth, simple curves; those that approach being spherical are called vesicles. Vesicles are quite common just below the desert pavement of soils in western Arizona. Channels are voids with a cylindrical form larger than those resulting from normal packing. Chambers are similar to vesicles except that they are interconnected with channels. Other voids may be planer in form, such as those associated with cracks and fissures.

FIGURE 7. Crystal Structure of Kaolinite Clay (after P. W. Birkland, 1974)

FIGURE 8. Representative Diagrams of Forms of Soil Aggregation

The atmosphere extends into the soil through voids. Soil atmosphere is a natural continuation of the atmosphere. Although soil air is similar in some respects to atmospheric air, it differs in others. Soil air often is saturated with water vapor, has higher levels of carbon dioxide and may have less oxygen. The levels of gases present depend upon the activity of microorganisms in the soil and the gaseous exchange between soil and atmosphere. Additions of leaf litter greatly

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stimulate microbial activity, resulting in oxygen depletion and carbon dioxide increase. Exchange of oxygen and carbon dioxide with the atmosphere is by diffusion, a process that is hindered if soil voids are small and few. If the voids are filled with water, oxygen cannot easily diffuse in and oxygen that is present is used up, producing anaerobic conditions. The absence of oxygen and the soil chemical changes that result from it combine to make the soil useful only to water-adapted vegetation.

Air and water in the soil are almost complementary, because if the soil is saturated with water, air is displaced. In a saturated soil nearly all voids are filled with water. If the soil drains and all water in the larger voids is removed, the water lost is gravitational water (Figure 9). About two days after flooding or a heavy rain, a freely drained soil loses its gravitational water and is at field capacity. In this state, considerable amounts of water remain in finer voids because of capillary attraction.

If the soil continues to lose moisture from capillary water reserves, plants will be unable to obtain enough water for transpiration. Wilting occurs and the plant does not recover. This is the permanent wilting point. Although by convention the permanent wilting point is considered to correspond to a soil moisture tension of 15 atmospheres, the point is not actually fixed and varies according to the soil and plant involved. Some desert plants, for example, can obtain water from the soil against strong capillary forces. The amount of water held in soil between field capacity and permanent wilting point is the available water capacity. It varies according to soil texture and structure (Figure 10).

Additional water can be obtained from soil in the laboratory. Soil can be brought to an air-dry state, but it is rather variable, depending upon the relative humidity of the atmosphere where the soil is kept. Water held by soil at temperatures up to oven dry, 105 C (221 F), is hygroscopic water and is mostly unavailable to plants.

Water moves in soils through the soil voids by saturated flow. Rain or irrigation water infiltrates at the soil surface and continues downward by gravity and by capillary forces. In unsaturated soils movement is restricted, taking place slowly in response to capillary forces. Water also can move as vapor from a warmed soil layer into a cooler layer where condensation occurs.

FIGURE 9. Form of Water in Soil

Soil water dissolves soluble constituents and as such will constitute the soil solution which is the medium whereby plants are supplied with nutrients. Inorganic salts dissociate into ions in solution, many of which are attracted to and adsorbed onto clay and humus particle surfaces in exchangeable form. An equilibrium may be approached between exchange position numbers and ions still in the soil solution. The concentration of hydrogen ions in solution is defined by the pH scale. Neutrality is pH 7; values below pH 7 are acid, above pH 7 alkaline. Most Arizona soils are between pH 6 and 8. A few sodic soils have pH values exceeding 9 and some leached soils in the higher elevations may have pH values as low as 4.

1.4. Soil Horizons

A unique feature of most soils is the horizontal layering that generally is seen in a vertical cut from the soil surface to the underlying rock or unconsolidated rock material. The layers are soil horizons and taken together they constitute the soil profile. Each horizon may differ from its neighbors in color and depth and in physical and chemical properties. Horizons are designated according to soil profile position and the processes that created them. Soil horizon notations currently used in the United States by the National Cooperative Soil Survey (Soil Survey Staff, 1981) are described below.

The capital letters O, A, E, B, C and R represent master soil horizons.

O horizons are dominated by organic material, except limnic layers that are organic. Some are saturated with water for long periods of time or were once saturated but are now artificially drained. Others have never been saturated. They may be subdivided into:

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A horizons are mineral horizons that have formed at the surface or below an O horizon. They are characterized by an accumulation of humified organic matter intimately mixed with the mineral fraction and not dominated by properties characteristic of E or B horizons, or have properties resulting from cultivation, pasturing or similar kinds of disturbances. In the latter case the horizon is designated Ap.

FIGURE 10. Water in an Unsaturated, Coarse-Textured Soil

E horizons are subsurface mineral horizons that are lighter in color because of removal of organic matter, iron and aluminum oxide minerals, silicate clay or some combination of these that leaves a concentration of sand and silt particles of quartz and other resistant minerals.

B horizons form below an A, E or O horizon and are dominated by the obliteration of all or much of the original rock structure and by illuvial concentration of silicate clay, iron, aluminum, humus, carbonates, gypsum or silica, above or in combination; evidence of removal of carbonates; residual concentration of sesquioxides; coatings of sesquioxides that make the horizon conspicuously lower in value, higher in chroma or redder in hue than overlying and underlying horizons without apparent illuviation of iron; alteration that forms silicate clay or liberates oxides or both and that forms granular, blocky or prismatic structure if volume changes accompany changes in moisture content; or any combination of the aforementioned. Subdivisions of B horizons that are common in Arizona soils include:

C horizons or layers, excluding hard bedrock, are little affected by pedogenic processes and lack properties of O, A, E or B horizons. Most are mineral layers, but limnic layers, whether organic or inorganic, are included. The material of C layers may be either like or unlike that from which the solum presumably formed. A C horizon may have been modified even if there is no evidence of pedogenesis. Subdivisions of C horizons in Arizona soils include:

R horizons are used to designate hard or very hard bedrock.

AB is a horizon transitional between A and B horizons dominated by properties characteristic of an overlying A horizon but having some subordinate properties of an underlying B horizon.

BA is a horizon transitional between A and B horizons dominated by properties characteristic of an underlying B horizon but having subordinate properties of an overlying A horizon.

EB is a horizon transitional between E and B horizons dominated by properties characteristic of an overlying E horizon but having some subordinate properties of an underlying B horizon.

BE is a horizon transitional between E and B horizons dominated by properties characteristic of an underlying B horizon but having some subordinate properties of an overlying E horizon.

BC is a horizon transitional between B and C horizons dominated by properties of an overlying B horizon but having some subordinate properties of an underlying C horizon.

A/B, B/A, E/B, B/E, A/C and B/C are horizons in which distinct parts have recognizable properties of two kinds of horizons. The two capital letters are separated by a virgule (/). The first letter is that of the horizon that makes up the greater volume.

Limnic layers are composed of materials deposited or formed in freshwater such as coprogenus and diatomatious earth and marl.

A hypothetical soil profile with many of the different kinds of horizons in Arizona is shown in Plate 15. No given soil would contain all these horizons.

1.5. Soil-forming Factors and Processes

Horizons in soil reflect the combined effects of soil-forming factors and soil-forming processes. Climate, organisms (mostly vegetation), parent material, topography and time are the five soil-forming factors (Figure 11). The sequence of changes that the parent material goes through in forming a specific soil are soil-forming processes. Additions to, removals from and vertical transfers and transformations within the soil are the four basic kinds of soil-forming processes. The relative importance and specific nature of these processes vary according to the effects on the soil of the soil-forming factors, which also affect the evolution of the soil and its profile (Figures 12 and 13).

1.5.1. Factors

The importance of the soil-forming factors has been recognized for about 100 years. More recently, Jenny ( 1941, 1961 and 1980) extended the concept and attempted to give them a rigorous mathematical treatment. Jenny's approach is to show genetic or geographical relationships among soils using a ‘‘fundamental’’ equation of state,

s = f (cl, o, r, p, t),

where s is any soil property, f indicates a functional relationship, cl is climate, o is biological activity (organisms), r is topography or relief, p is the soil parent material and t is the time during which the soil forms. Jenny also attempted to

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evaluate the influence of each factor on different soil properties by keeping the other factors constant.

FIGURE 11. Schematic of Various Factors That Affect Soil

For example, to assess the role that precipitation (cl) has on a soil property, the soil property is measured under different precipitation conditions under the same or similar temperature conditions, topography and parent materials, and using soils of similar age. Although practical difficulties are encountered in applying this equation, some very useful generalizations have been obtained. Because of the importance of these soil-forming factors, their roles in the formation of Arizona soils will be discussed briefly.

Climate. Climate includes type and amount of precipitation, temperature, humidity, evapotranspiration, duration of sunshine and a number of other variables.

Precipitation and temperature are the main parameters of climate and are important in influencing the nature of the soil that forms. Precipitation is important because of its effect on the soil-moisture regime. Much of the precipitation that falls in Arizona is lost to runoff and evaporation from vegetation and soil surfaces. Still more of the moisture that enters the soil is taken up by vegetation through roots and transpired from plant leaf surfaces, further reducing soil moisture levels. In low precipitation areas, not enough water is added

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to the soil for through leaching. Water that enters the soil penetrates to such a limited depth that soluble constituents are not removed from the soil profile. Thus, most of Arizona that does not receive much precipitation contains soils that have soluble salts and carbonates, and generally are neutral to alkaline in reaction (pH). The higher elevations in Arizona receive more precipitation than the lower, and more water percolates through the soil to remove soluble substances. Soils in these areas generally are leached of soluble salts, lack carbonates and are more acid.

FIGURE 12. Representative Illustration of the Evolution of Soil (after E. M. Bridges, 1978)

Lack of moisture in the lower, arid regions of the state also limits vegetative production and inhibits soil organic matter buildup. Soils in the higher, more humid regions, on the other hand, are higher in organic matter because of greater vegetative growth.

Temperature is important in soil formation because it influences the effectiveness of precipitation. Arizona has two principal rainy seasons, December through March and July through August. Higher temperatures associated with summer rains cause much more moisture loss by evaporation and transpiration than is lost during the winter season; therefore, soil moisture levels and leaching rates are much less than during the winter season. Soils at the higher elevations in Arizona that receive more precipitation also are cooler. Thus, the amount of water available to moisten soils is greater and leaching intensities are enhanced further.

Temperature also influences biological activity since greater vegetative production occurs in soils with higher temperatures. But in Arizona limited moisture restricts vegetative growth in the warmer areas and accelerates decomposition of the organic material available to the soil. Decomposition of soil organic matter in the more humid, cooler and higher elevation soils is slower and the buildup of organic matter is higher.

The combined effect on soil properties of increasing precipitation and decreasing temperatures as elevations increase has been studied in many areas of the world. Two such studies were made in Arizona by Martin and Fletcher ( 1943) in the Pinaleno Mountains near Safford, and by Whittaker et al ( 1968) in the Santa Catalina Mountains near Tucson. The latter authors determined that soil organic matter increased by about 3.23 percent per 1,000 m (3,300 ft) increase in elevation up to 2,000 m (6,600 ft). A higher rate of increase was observed in the coniferous forests above 2,000 m (6,600 ft). Surface soil pH in the Santa Catalina Mountains decreased with increasing elevations by 1.29 units per 1,000 m (3,300 ft), according to Whittaker et al ( 1968).

Organisms. Organisms, particularly vegetation, long have been recognized as important factors in soil genesis. Vegetation is controlled or strongly influenced by climate, however, which makes a separate evaluation of its role in soil formation difficult. Jenny ( 1958, 1980) and Crocker ( 1960) have described theoretical methods of dealing with this problem.

Plants may influence soil properties in a number of ways. The amount and depth of organic matter is related to the type of vegetation growing on the soil. Forest soils tend to have organic matter concentrated at or near the surface because most organic matter added to the surface is leaf litter. Grass vegetation, on the other hand, generally contributes annually much more organic matter to the soil and to greater depths below the surface; the total amount of organic matter throughout grassland soils often is higher than that in forests. Forest soils generally are more acid and more strongly leached than grassland soils. Although these differences may be attributed mostly to the climatic conditions associated with the vegetation, the differing chemical nature of organic matter formed during the decomposition of the two kinds of plant materials also may contribute.

Zinke ( 1962) and Zinke and Crocker ( 1962) described and applied a method of assessing the effects of individual trees on soil properties. Since the life of a tree is rather short from a soil-formation point of view, only soil properties that change rapidly in response to the living tree would be affected.

Soils near the tree trunk were compared with those under the tree canopy and in adjacent open areas or under a neighboring tree. Soils near the trunk are influenced by bark litter and rain diverted by the tree as stem flow. Soils under the canopy are influenced by leaf litter and canopy drip. Soils in the open are much less affected by the tree. The general trend for most tree species observed by Zinke ( 1962) was for soil pH to increase radially from the trunk outward. Similarly, nitrogen content, exchangeable cations and cation exchange capacities were low near the trunk, increased to a maximum under the canopy some distance from the trunk and then generally declined farther outward.

The approach of Zinke ( 1962) was applied by Tiedemann and Klemmedson ( 1972, 1973, 1977) and Barth and Klemmedson ( 1978, 1982) to mesquite and paloverde on the Santa Rita Experimental Range south of Tucson. Tiedemann and Klemmedson ( 1972, 1973, 1977) found more nitrogen, potassium, sulfur, soluble salts and organic matter in soils under mesquite trees than in open spaces between trees. Mesquite trees redistribute nutrients that are absorbed through roots to a zone mostly beneath the canopy where most leaves and other plant parts accumulate.

FIGURE 13. Generalized Soil Profile (after D. Hillel, 1982)

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Barth and Klemmedson ( 1978) observed that the surface organic litter derived from both velvet mesquite and paloverde decreased in weight as distance from the plants' center increased, whereas litter originating from understory vegetation displayed weak spatial patterns in dry weight. They also found that the percent of soil nitrogen and organic carbon decreased with horizontal distance from the plants' center and with depth, to 60 cm (24 in).

Soil pH tended to increase under velvet mesquite but to decrease with horizontal distance from the center of paloverde plants. Soil pH values increased with depth under both velvet mesquite and paloverde.

These distribution patterns of soil nitrogen, organic carbon and pH are most marked in the surface soil (0 to 5 cm, or 0 to 2 in). Barth and Klemmedson ( 1982) further found no apparent increase in soil nitrogen and organic carbon with increased size of paloverde plants, whereas under velvet mesquite they estimated that the soil accumulated nitrogen at the rate of 11.2 g/m2 (.036 oz/ft2) per meter (39 in) of height and organic carbon at the rate of 0.11 kg/m2 (3.6 oz/ft2) per meter (39 in) of height.

Topography (or Relief). Soil properties commonly vary laterally with topography. One cause is the steepness of slope and its influence on surface-water runoff and erosion. Steeper slopes generally have rapid runoff and less moisture entering the soil. Additional moisture may be received by soils in low-lying areas from upslope runoff.

Erosion potential is greater on steep slopes. Surface soil material is removed more rapidly from them than from more gentle slopes. Consequently, soils on steep slopes often are shallow and lack horizon development. The eroded material usually is deposited as alluvium in low-lying areas, along stream channels or in floodplains. This alluvium then serves as parent material for soils that form in those areas.

Low areas sometimes have high water tables and the soils are saturated with water. Saturated soils may lack oxygen, and chemical reactions favored in anaerobic conditions occur, such as reduction of iron and manganese. High amounts of soluble salts accumulate readily in saturated soils because water containing dissolved salts moves toward the soil surface by capillary rise. When this water evaporates from the soil the soluble salts that it carried remain in the soil. Over time, fairly large accumulations of soluble salts accrue by this process and the soil will support growth only of salt-tolerant plants.

Another aspect of topography that affects soil properties is hillslope orientation, mostly because of its influence on the microclimate. Soil temperature regimes are not alike on slopes of different exposure because of disparate solar radiation, soil moisture, vegetative cover and wind direction and velocity (Figure 14). Southern sun-facing slopes usually receive more heat from solar radiation and are drier than northern slopes. Different soils, therefore, form on sun-facing slopes than do on shady slopes. Many hills and low mountains in the intermediate to higher elevations in Arizona have distinctly different vegetation on north- and south-facing slopes. South-facing slopes commonly are covered with grass, the north-facing slopes with woodland or forests.

FIGURE 14. Effects of Slope and Aspect on Soils (after E. M. Bridges, 1978)

The soils also are different. For example, Green's Peak, a cinder cone in southern Apache County, has an Engelmann spruce forest--with some admixture of quaking aspen--on the north-facing slope and a grassland on the south-facing slope (Figure 15). Hendricks and Davis ( 1979) reported that forested soils are more acid, higher in clay and free iron oxides content, and have lower nitrogen content than grassland soils. Although slopes that face the wind receive more precipitation than those that do not, the wind also has a drying effect and may carry fallen snow off the slope.

Parent Material. Jenny ( 1941) defined parent material as being the initial state of soil systems. When parent material is exposed to environment at a site, soil-forming processes begin.

Parent material exerts the greatest influence in arid regions and during the early stages of soil formation. In humid regions and in older, more extensively weathered and evolved soils parent material differences tend to be less evident.

Parent material influences development of many soil properties to varying degrees. Soil constituents such as sand-and silt-sized particles may be inherited directly from the parent material with little if any change. Other constituents such as clay minerals evolve through chemical changes in material components. Parent material influences may or may not be obvious. For example, coarse-grained parent materials often produce coarse-textured soils. Coarse-grained parent

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materials that weather easily to form clay, on the other hand, tend to produce fine-textured soils. Thus, coarse-grained granitic rocks composed mostly of minerals not susceptible to chemical weathering generally produce coarse-textured soils high in sand. Basaltic rocks that contain mostly minerals quite susceptible to chemical weathering yield fine-textured soils high in clay content. Soils formed on granitic rocks also generally are deeper than soils formed from basalt because granitic parent materials are more susceptible to physical disintegration than basaltic parent materials in which smaller mineral grains are tightly interlocked.

FIGURE 15. Green's Peak

Unconsolidated parent material such as stream-deposited sediments in Arizona valleys may influence soil formation through the ability of its texture to retain or otherwise affect moisture movement. Sediment texture influences the rate and depth of leaching, which, in turn affects many soil properties, including the zone of carbonate accumulation as illustrated in Figure 16.

FIGURE 16. Depth of Carbonate Accumulation in Soils Relative to Parent Material Texture

Because of chemical makeup some parent materials produce soils that have higher inherent fertility than others and that support greater vegetative growth. Fertility and resultant vegetation produce soils richer in organic matter. Welch and Klemmedson ( 1973), for example, reported that in the Colorado Plateau ponderosa pine zone near Flagstaff basic volcanic rocks (andesite and basalt) produced soils in both tree and in grass ecosystems that contained more nitrogen than soils derived from acid volcanic rocks (rhyolite).

Time. The length of time that a soil has been forming dictates the degree of horizon expression and other pedogenic soil features that evolve given a combination of

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climate, organisms, topography (or relief) and parent material. The time necessary for various features to evolve depends not only on the four other soil-forming factors, but on the nature of the particular feature that is formed. Formation of zones of soluble salt accumulation and A horizons by organic matter buildup happens relatively rapidly, less than 100 to 1,000 years (Yaalon, 1971). Other features, such as argillic (clay enriched) and petrocalcic (lime-cemented pan) horizons, form quite slowly, 1,000 to 10,000 years or more.

Most soil properties develop more rapidly during the initial stages of soil formation. Decreasing rates of development follow until a steady-state eventually is reached. Steady-state is a condition in which no apparent additional change in a soil property occurs because of time. Thus, rates of soil formation initially are rapid and then slow until it is zero, as demonstrated by chronosequence studies reviewed by Stevens and Walker ( 1970). However, a more recent review by Bockheim ( 1980) cast some doubts about soils actually reaching a steady-state with environments.

Several approaches have been made to assess the role of time in soil formation:

  • by determining the relative stage in soil formation from a time zero, the estimated time when parent material was first exposed to weathering;
  • by determining the rate of formation of a unit depth of soil or soil horizon;
  • by referring to the age of the slope or landform on which the soil formed; and
  • by absolute dating (radiometric dating) of a part of the soil profile (Bunting, 1965).

1.5.2. Processes

Many possible physical, chemical and biological reactions and interactions occur in soils that are categorized as soil-forming processes. These numerous processes are grouped into four basic kinds of changes: additions to, removals (losses) from, vertical transfers and transformations within the soil (Simonson, 1959, 1978). Vertical transfers and transformations are grouped once again under the heading of Translocations in this discussion. Acting under the influence of and in conjunction with the soil-forming factors, these processes produce the variety of soils that exist.

Additions. Organic matter that accumulates in soils includes the remains of organisms, plants and animals, but mostly plants. When plant debris falls onto the soil surface it may accumulate to form a surface litter layer or it may be mixed into the surface mineral horizons. Organic matter also may be added directly to the soil from roots and by animals.

Dissolved substances and solid particles are added to soil by precipitation. Solid particles also are added by aeolian action. Aeolian materials can have a strong influence on soil properties. Slow carbonate accretion in soil often is attributed to aeolian actions. Over a long period of time carbonates build up to relatively high levels in some soils, although the soil parent material lacks carbonates and is low in other calcium-containing minerals. The buildup of soluble salts in some other Arizona soils by this mechanism also is recognized (Nettleton et al, 1975).

Soils in floodplains and that are subject to flooding receive fresh deposits of alluvium from the floodwaters. Floodplain soils are subject to more frequent flooding and lack horizons because not enough time passes for their development before another flood deposits new material on the soil surface. The soils of floodplains, then, are characterized by limited pedogenic horizon development, but do have layers of material deposited by different flooding events.

Losses. Precipitation sometimes exceeds evapotranspiration at higher elevations in Arizona. Where this occurs downward percolating water dissolves readily soluble salts that later are carried from the soil by underground water into an aquifer or stream channel. Over a long time, percolating water removes materials that are less soluble, such as calcium carbonate minerals. The soils at higher elevations on the Kaibab Plateau north of the Grand Canyon, for instance, are free of calcium carbonate even though they formed from Kaibab limestone. The calcium carbonate originally in Kaibab limestone leached away and left noncarbonate impurities. These and added aeolian materials constituted the parent material of the soils. Leaching processes also may remove exchangeable bases such as calcium, magnesium and potassium and leave behind hydrogen or aluminum. Thus, a major effect of leaching is gradually to make the soil more acid.

Losses of soil material by wind and water erosion are common. Wind erosion is especially active in the more arid regions of Arizona where the vegetation is sparse and much of the surface is exposed. In these areas, wind may remove smaller soil particles and leave desert pavement, a stone or gravel layer on the soil surface. Desert pavement protects additional small particles underneath it from further wind erosion. Soils formed in some floodplains from finer alluvium and those formed from fine-grained sedimentary rocks that do not contain gravel may continue to be subject to wind erosion unless vegetation is established, either naturally or by man.

Water erodes soil material by raindrop splash and surface runoff, often called sheet erosion. Sheet erosion removes smaller soil particles, organic matter and soluble constituents from the soil surface. Rill and gully erosion results when surface runoff concentrates in surface depressions and forms well-defined channels, or rills. As runoff continues to erode the rills, gullies form and water removes soils from lower horizons as well as from the surface. Soils unprotected by vegetation or by desert pavement are most susceptible to water erosion. The high-intensity rains characteristic of the summer thunderstorms in Arizona are most effective in producing voluminous, rapid runoff that carries away massive amounts of soil.

Translocations. The movement of dissolved materials and very small particles from one part of the soil to another is very important in the formation of soil horizons. The soluble salts are removed from other soils and/or weathering rocks, especially salt-rich sediments, and carried by groundwater. Dissolved salts introduced to soils that have high water tables by groundwater are drawn upwards into the soil by capillary action and deposited as the water evaporates. As a result, these soils contain concentrations of salt at or near the surface that, in some cases, can develop as a surface encrustation.

Dissolved materials are carried downward in well-drained soils by percolating water, unless changing chemical conditions or dehydration cause them to precipitate. This is the

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principal mechanism by which carbonate-enriched horizons form in soils. Calcium carbonate in the upper part of the soil dissolves in water and is carried downward by it to be precipitated at a lower depth. In addition, calcium released by chemical weathering of noncarbonate minerals also percolates down through the soil to be precipitated as carbonates.

The zone where the carbonates are deposited represents the average depth of penetration of the percolating water. Given enough time, (usually more than 10,000 years) this process may form a horizon completely plugged, cemented and indurated with carbonates. Most soils in the arid and semiarid regions of Arizona have been affected by this process. At the higher elevations in Arizona, leaching removes carbonates from the profiles of the more humid soils.

Soil material also translocates as small particles in suspension from the surface horizons by water percolating to lower horizons in the soil. This process involves dispersion of clay and other colloidal soil materials and their movement downward with soil water, and finally by their deposition in lower horizons. For dispersion to take place, the upper horizons must be free of carbonates and low in soluble salts, and have a pH above about 5. Once dispersed, clay particles move through the larger voids by noncapillary or gravitational water, according to most theories. The suspended clay particles may be flocculated by calcium and magnesium from carbonates or by soluble salts in lower horizons. The suspended clay also is deposited when percolating water reaches a drier subsoil, causing the noncapillary flow to cease. When this happens, clay is deposited on the walls of the larger voids as water withdraws.

One of the most important results of clay translocation is development of a Bt horizon enriched in clay, a horizon referred to as being argillic. Clay content in the argillic horizon can be increased considerably compared with the amount remaining in overlying A or E horizons. Thus, a common feature of soils that undergo clay translocation is development of films or coatings of clay lining walls of larger voids and on ped surfaces. These coatings in argillic horizons sometimes are destroyed in Arizona soils by the shrinking and swelling of the soil when drying and wetting cycles occur.

Most soils swell when wet and shrink as they dry. In some soils the magnitude of change with wetting and drying may be quite small, whereas in other soils it may be fairly large, for instance deep, wide cracks may extend about 1 m (3.3 ft) into some soils during dry seasons. Soil material from the surface often falls into the cracks. When the soil becomes wet and swells, the cracks close exerting pressures upward because of the added excess material that fell into the lower horizons (Figure 17). Soil material then is forced toward the surface. With time, these soils invert themselves. Material originally in the subsoil moves up and becomes the surface soil and the surface soil becomes the subsoil. Soils that exhibit this phenomenon are Vertisols. Arizona Vertisols formed mostly from basalt in the central and east-central part of the state at elevations ranging from 1,215 to 1,970 m (4,000 to 6,500 ft). A common feature of Vertisols in Arizona is the large accumulation of basalt stones on the soil surface. Soil body stone content is low, suggesting that the stones originally in the soil worked up to the surface over time. Here they accumulate since they are too large to fall into cracks during the dry season.

Nutrient elements absorbed by plant roots at some depth within the soil move upward through the plant in its tissue. When the plant dies or sheds leaves, these nutrient elements are returned to the soil. This process, nutrient cycling, transfers elements from various horizons within the soil to the surface. Some large trees extend roots several meters into the soil and, sometimes, into cracks in underlying rocks where they remove nutrients later deposited on the surface.

Soil animals play an important role in moving solid particles from one part of the soil to another. These animals' activities, whether on a small scale such as those of insects, worms and other invertebrates or on a larger scale by burrowing snakes, gophers and mice, mix soil material from one horizon with that of another.

FIGURE 17. Churning Action of Vertisols

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FIGURE 18. Selected Recreation Areas in Arizona (after M. E. Hecht and R. W. Reeves, 1981)

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