CHARACTERISTICS OF KARST AQUIFERS IN TENNESSEE
Karst aquifers contain a variety of flow regimes, ranging from rapid turbulent flow in freely draining conduits to slow laminar flow through bedrock-fracture networks, cave and fracture sediments, or regolith. The relative importance of different types of flow varies with depth, horizontal location, short-term changes in hydrologic conditions, and long-term geologic evolution of the aquifer. The spatial and temporal distribution of different flow regimes largely defines the overall hydraulic character of a karst aquifer.
Under the best of circumstances, chlorinated solvents are difficult to find and recover once they have entered the subsurface (Cherry, 1996). The distinctive hydraulic characteristics of karst aquifers compound this difficulty. Sampling strategies, analytical techniques, and approaches to monitoring that work in unconsolidated granular or fractured rock aquifers commonly give erroneous results in karst (Quinlan, 1989, 1994; Barner and Uhlman, 1995).
Karst conduits provide pathways for the rapid transport of water and contaminants. The presence of conduits, natural openings formed by the chemical dissolution of limestone or other soluble rock, often confounds efforts to understand or model karst aquifers through standard analytical techniques based on Darcy's Law (White, 1988; Ford and Williams, 1989). The most familiar examples of karst conduits are large explorable caves with diameters on the order of meters to tens or hundreds of meters (Barr, 1961; Matthews, 1971; White and White, 1989). Most karst conduits are much smaller than explorable caves, with widths ranging down to about a few millimeters (White, 1988; Ford and Williams, 1989). Conduit shape and orientation are extremely variable. Some conduits are nearly circular in cross section, but most are long and thin, following the orientations of bedding planes, joints, and similar sheet-like openings. Conduits a few centimeters across can dominate ground-water flow in an aquifer, even though they account for only a small part of water-bearing pore space (Quinlan and Ewers, 1985; Smart and Hobbs, 1986).
Diffuse-flow networks in karst aquifers include integrated fracture systems in bedrock and intergranular pores in regolith and unconsolidated cave fill. The size, distribution, tortuosity, and interconnection of voids in fractured rock or unconsolidated material determine how readily water will move through diffuse-flow systems. Fractured rock and unconsolidated materials transmit water much more slowly than conduits but may store large quantities of water. Fracture networks and unconsolidated materials typically represent a much greater proportion of aquifer volume than do large conduits. Much of the flow carried by conduits first passes through diffuse-flow networks. Chemical and isotopic analyses of water seeping into a cave in Israel showed that water was stored in the regolith and small fractures for periods ranging from days to decades before entering the karst-conduit system (Even and others, 1986). Bedrock fractures, regolith macropores, and intergranular pores of unconsolidated material also provide potential storage for contaminants.
The behavior and ultimate fate of a given contaminant mass in a karst setting will reflect the relative importance and detailed geometry of the conduits and diffuse-flow zones into which the contaminant has been introduced and the degree of hydraulic coupling between the flow systems. Contaminants introduced into sinkholes may move directly into connected conduits. More typically, contaminants accumulate in diffuse-flow zones within karst aquifers. In such cases, the hydraulic efficiency of the diffuse-flow network and the degree of integration between the diffuse-flow and conduit components of the aquifer exert a major influence on the residence time and delivery of dissolved-phase contamination into the conduit system and, eventually, to a discharge point.
The concepts presented by White (1969, 1977), Quinlan (1978), Stringfield and others (1979), Chen (1988), Gunn (1992), and Quinlan and others (1992) provide a basis for identifying major factors that control the recharge, flow, and discharge characteristics of karst aquifers in a given climatic setting. These factors include geologic factors such as lithology, geologic structure, and stratigraphy, and geomorphic factors such as topographic relief, landforms such as sinkholes and caves, and the thickness and properties of the regolith.
Lithology is probably the single most important determinant of whether karstification occurs in a given climatic and tectonic setting. In the Eastern United States, some degree of karst development can be assumed in any rocks that contain a significant proportion of carbonate minerals (Quinlan, 1978; Quinlan and others, 1992). Karst development is generally proportional to carbonate or other soluble mineral content, mechanical strength, and absence of primary porosity (Sweeting, 1972; White, 1988; Ford and Williams, 1989). Carbonate rocks with high primary porosity or high proportions of insoluble minerals will generally have lower levels of karst development. However, such rocks are likely to have some degree of dissolution enlargement of joints, bedding planes, and other voids and may include areas with considerable karst development (Brahana and others, 1988; Wolfe, 1996a, b).
Geologic structure and stratigraphy determine the spatial arrangement of karst and nonkarst rocks. Regional fold systems, such as that in the Valley and Ridge Physiographic Province (Fenneman, 1938), tend to constrain the direction of ground-water flow, typically down-dip or along strike (Bailey and Lee, 1991). The numerous faults that commonly accompany regional folding may offset confining units and allow hydraulic connection between stratigraphically separate karst aquifers.
In flat-lying or gently dipping strata, the thicknesses and relative positions of karst and nonkarst rocks are major controls of the overall ground-water flow regime and of the movement and storage of contaminants. The gently dipping carbonate strata of Middle Tennessee form a series of alternating karst aquifers and clay-rich leaky carbonate confining units. Thick karst aquifers are important sources of ground water. These aquifers may represent high potential for deep infiltration of low-viscosity DNAPL's such as chlorinated solvents.
Topographic relief commonly reflects lithology and geologic structure, but its effects on karst development and the hydraulic response of karst aquifers can be considered independently. The difference in elevation between the top of a karst aquifer and regional base level is the maximum hydraulic head driving water through the aquifer (White, 1977). The greater that driving force, the greater the potential for karst development, when geologic and climatic factors are equal. White (1977) explicitly incorporated topographic relief relative to base level into his karst classification. Similarly, Quinlan and others (1992) recognize that the distribution of dissolution porosity relative to regional base level is a critical determinant of the volume of water that remains in long-term storage in a karst aquifer. If most of the dissolution porosity in a karst aquifer lies above regional base level, flow through the aquifer will be relatively rapid and storage minimal (White, 1977; Quinlan and others, 1992). A karst aquifer with most of its dissolution porosity below regional base level will have a correspondingly higher volume of long-term storage and, for a given conduit geometry, slower flow rates. Regional base level can rise or fall relative to the location of secondary porosity during the geologic lifespan of a karst aquifer, causing adjustments in flow and storage characteristics.
Karst landforms are critical factors in routing surface runoff and recharge. Surficial karst landforms in the humid, unglaciated parts of the United States include simple and compound sinkholes (uvalas), residual hills (knobs), blind valleys, disappearing streams, vertical shafts, cave openings, springs, and seeps (White and others, 1970; White, 1988, 1990). With a few exceptions, most of these landforms were formed by, and continue to mediate, hydraulic connections between the surface and subsurface; they are points of ground-water recharge, discharge, or, intermittently, both. The distribution of sinkholes, shafts, disappearing streams, and related karst depressions is a critical determinant of the rapidity of ground-water recharge and the hydraulic response of an aquifer to rainfall events. Such depressions also provide efficient pathways for contaminants to enter karst aquifers.
Regolith and landforms are important components of karst aquifers because they largely control the rate, depth, and distribution of recharge (White, 1977, 1988; Quinlan and others, 1992). These factors also influence movement of contaminants from the surface to underlying karst aquifers. Regolith includes soil, residuum, colluvium, and alluvium. In humid areas, such as Tennessee, where bedrock outcrops compose a small part of the land surface, contaminants released at the surface must pass through the regolith to reach an underlying karst aquifer.
One important control on the movement of fluids through regolith is macro-porosity. Macropores include cracks, pipes, root channels, animal burrows, and other visually discernible openings with dimensions on the order of 0.1 cm to several centimeters or larger (Germann and Beven, 1981). Such openings are present in all soils and in most natural unconsolidated materials. A wide variety of processes including plant growth, animal activity, and shrinking and swelling in response to changes in temperature or moisture content form macropores (Germann and Beven, 1981).
Karst processes greatly increase the likelihood of macropore formation in several ways. The gradual dissolution and lowering of the bedrock surface generally occurs unevenly (White, 1988; Ford and Williams, 1989) so that the regolith is subjected to slowly but constantly changing stresses. More episodic karst processes such as sinkhole collapse leave large voids in the soil surface that are directly connected to karst conduits (White, 1988). Once a collapse occurs, tension or shear cracks may develop as the surrounding soil adjusts to the collapse through mass movements (Kemmerly, 1981). Dissolution openings may be retained as relict soil structures in carbonate-rock residuum. Numerous studies in Tennessee and adjacent states have documented rapid, preferential infiltration and routing of water through soil macropores (Thomas and Phillips, 1979; Watson and Luxmoore, 1986; Quinlan and Aley, 1987; Wolfe, 1996a, b).
The interface between bedrock and regolith commonly is highly irregular in karst settings. Lithologic and structural variation and numerous other factors typically produce a complex top-of-rock topography characterized by bedrock pinnacles, pits, channels, and weathered rock fragments mixed with finer grained regolith. This complicated mixture of consolidated and unconsolidated material, known as epikarst or the subcutaneous zone (Williams, 1985), is one of the least understood aspects of karst. Epikarst can store infiltrating ground water (or contaminants) or provide a direct route to the underlying aquifer (White, 1988; Ford and Williams, 1989). In some settings, epikarst forms an important, even locally dominant, aquifer (Haugh and Mahoney, 1994; Julian and Young, 1995). Quinlan and others (1992) regard the presence of epikarst as a given in carbonate settings unless evidence is provided for its absence.
Unconsolidated materials affect the movement of water and contaminants within karst aquifers below the epikarst zone (Gale, 1984). Dissolution enlargement of voids commonly leaves insoluble residue from the carbonate rock in caves, joints, and bedding planes. Such residual fill can remain in place, where it interferes with flow, or it can be mobilized as cave sediment. Infiltrating rainwater can carry the finer fraction of unconsolidated material from the regolith downward into karst openings in the bedrock. Disappearing surface streams also transport sediment into karst aquifers (White, 1988; Ford and Williams, 1989).
Carbonate rocks underlie most of Middle Tennessee and large areas of East Tennessee. The carbonate areas of Tennessee can be divided into regions based on geologic structure, stratigraphy, relief, regolith thickness, and karst landforms. Six karst regions will be considered in this report: (1) the inner Central Basin, (2) the outer Central Basin, (3) the Highland Rim, (4) coves and escarpments of the Cumberland Plateau, (5) the Valley and Ridge, and (6) the western toe of the Blue Ridge (fig. 5). The delineation of the six karst regions was based on their general physiographic and hydrogeologic characteristics (table 3) and selected geomorphic criteria such as depth to bedrock, land-surface slope, streamflow recession characteristics, and sinkhole density (table 4).
The Central Basin ("Nashville Basin," Fenneman, 1938) is an elliptical topographic basin trending northeast to southwest through the middle of Tennessee (fig. 5). Most of the basin is underlain by carbonate rocks of Ordovician age that were exposed by weathering and erosion of the central part of the Nashville (structural) Dome (Fenneman, 1938; Wilson, 1962; Miller, 1974). These carbonate rocks include several relatively pure limestones with well-developed karst conduit systems. The relatively pure limestones are separated by shaley confining units (Newcome, 1958). At the base of the Ordovician carbonate sequence is the Cambrian-Ordovician Knox Group (Newcome, 1958). In the Central Basin, the upper contact of the Knox Group is an erosional unconformity, generally 100 to 300 m below land surface. The base of the Knox Group is approximately 1,600 to 2,000 m below land surface (Newcome, 1958; Newcome and Smith, 1962). Newcome (1958) considered the paleokarst of the upper 30 m of the Knox Group to be a reliable aquifer in the Central Basin but noted variable water quality and generally low well yields.
The Central Basin can be subdivided into two concentric zones: a relatively flat inner basin and a more hilly outer basin. For this report, the boundary between the inner and outer Central Basin (fig. 5) corresponds roughly with the outcrop of the contact between the Hermitage Formation and the Carters Limestone. In comparison with the outer Central Basin, the inner Central Basin has thinner soils, higher sinkhole density, lower cave density, and lower land-surface slope (tables 3 and 4). The difference in cave density reflects the greater topographic relief of the outer Basin rather than a difference in karst development. Large (greater than 1 m) conduits are more likely to be explored and designated as caves when they drain freely and have readily accessible openings. Both of these conditions require a certain amount of topographic relief and are more commonly met in the hillier outer Central Basin than in the flatter inner Basin.
The most productive aquifers in the inner Central Basin are zones of conduit flow in the Carters, Ridley, and Murfreesboro Limestones (fig. 6), of which the Ridley Limestone is the most reliable (Newcome, 1958). Dissolution-enlarged conduits in these units typically are concentrated within 50 m of the land surface (E.F. Hollyday, U.S. Geological Survey, oral commun., 1996). Large explorable caves, such as Snail Shell Cave developed in the Ridley Limestone (Barr, 1961), are present, but most karst conduits are a few centimeters or smaller in width (J.V. Brahana, U.S. Geological Survey, oral commun., 1996). Important potential confining units include the Lebanon and Pierce Limestones (Newcome, 1958).
The capacity of a geologic unit to develop and maintain open conduits is inversely related to its insoluble mineral content (Ford and Williams, 1989, p. 30). White (1974) reports chemical analyses from three rock cores from the Central Basin. Fourteen samples of the Carters Limestone, 14 samples of the Ridley Limestone, and 55 samples of the Murfreesboro Limestone had average insoluble residue contents of 9 to 10 percent. In contrast, 16 samples of the Lebanon Limestone and 3 samples of the Pierce Limestone had average insoluble residue contents of 24 and 14 percent, respectively (White, 1974).
The outer Central Basin has hillier topography than the inner Basin, reflected in a median land-surface slope of 2 degrees compared with 0 degrees for the inner Basin (table 4). Major rock units include the Leipers and Catheys Formations, the Bigby and Cannon Limestones, and the Hermitage Formation, all of Ordovician age (Newcome, 1958; Wilson, 1962). Dissolution-enlarged conduits in the Bigby and Cannon Limestones compose the most important aquifer in the outer Central Basin (Newcome, 1958). Zones of well-developed conduits in the Bigby and Cannon Limestones are typically concentrated within 30 m of the land surface (E.F. Hollyday, U.S. Geological Survey, oral commun., 1996; fig. 7). Reported values for insoluble residue in the Bigby and Cannon Limestones average less than 10 percent (Sprinkle, 1973; White, 1974). The Hermitage Formation is an important confining unit across much of the outer Central Basin (Newcome, 1958). Six samples from the Hermitage Formation analyzed by White (1974) had insoluble-mineral contents ranging between 14 and 82 percent and averaging 48 percent. The Leipers and Catheys Formations are less effective confining units than the Hermitage Formation and function locally as relatively low-yielding aquifers (Newcome, 1958). Insoluble residue contents in the Leipers and Catheys Formations are intermediate between those reported for the Bigby and Cannon Limestones and the Hermitage Formation (Smith, 1972; Sprinkle, 1973; White, 1974).
The carbonate sequence that underlies the Central Basin is more than 1,600 m thick (Newcome, 1958; Newcome and Smith, 1962). Karst development in many of these rocks provides substantial potential for the downward movement of low-viscosity DNAPL's such as chlorinated solvents. The high density of sinkholes, especially in the inner Basin, means that contaminants have efficient pathways to the subsurface. Even where sinkholes are absent, the thin soils of the Central Basin offer little or no resistance to infiltration of contaminants above the bedrock surface.
The numerous confining units in the Ordovician sequence have often frustrated developers, home owners, and well drillers by their ability to interfere with the vertical movement of economically significant quantities of water. However, the effectiveness of these confining units with respect to water is highly variable--both between units and within a given unit at different locations (Piper, 1932; Newcome, 1958). Even relatively effective confining units such as the Hermitage Formation transmit water freely at some locations because of fractures or lithologic variation (E.F. Hollyday, U.S. Geological Survey, oral commun., 1996). The confining units of the Central Basin are likely to be less effective as barriers to the downward movement of low-viscosity DNAPL's than to infiltration by water. Even where confining units stop or retard DNAPL movement, abandoned wells, commonly cased only a few feet near the surface, may provide localized but very efficient pathways for DNAPL to move from higher to lower stratigraphic units (Crawford and Ulmer, 1994, p. 46-48).
The Highland Rim is an undulating plateau that surrounds the Nashville Basin (fig. 5). The Rim is bounded on the east by the Cumberland Plateau, on the west by the Western Valley of the Tennessee River, and extends north and south into Kentucky and Alabama, respectively. Traditionally, the Highland Rim in Tennessee has been divided into eastern and western subdivisions (Miller, 1974). West of the Central Basin, the Highland Rim can be further subdivided into a smaller section (the Pennyroyal Plateau) north of the Cumberland River and a larger section south of the Cumberland River (Kemmerly, 1980; Smalley, 1980).
The entire Highland Rim in Tennessee is underlain by carbonate rocks of Mississippian age (Fenneman, 1938; Miller, 1974). From youngest to oldest, major rock units include the Ste. Genevieve Limestone (Pennyroyal Plateau only), the Monteagle Limestone (stratigraphic equivalent of Ste. Genevieve Limestone in the Eastern Highland Rim), the St. Louis Limestone, the Warsaw Limestone, and the Fort Payne Formation. The Ste. Genevieve, Monteagle, and St. Louis Limestones are relatively pure, mechanically strong limestones that develop extensive cave and sinkhole systems (Piper, 1932; Kemmerly, 1980; Mills and Starnes, 1983). The Warsaw Limestone is a variable unit consisting of thickly to thinly bedded limestone of variable purity with interbedded calcareous shales and sandstones (Piper, 1932). The Fort Payne Formation is even more variable, consisting of "an extremely heterogeneous and variable assemblage of siliceous and calcareous shale and sandy, cherty, and earthy limestone" (Piper, 1932). The base of the Mississippian sequence, and of the Highland Rim aquifer system, which encompasses this sequence (Brahana and Bradley, 1986), is the Upper Devonian to Mississippian Chattanooga Shale (Piper, 1932; Miller, 1974).
Three distinct topographic styles are characteristic of the Highland Rim as a whole and recur in varying proportions throughout the Rim's extent:
Topography, especially the distribution of sinkholes and barrens, is strongly correlated with bedrock geology (fig. 8). Sinkholes are most common where the greatest thicknesses of relatively pure limestones such as the Ste. Genevieve and St. Louis Limestones crop out. Barrens are concentrated in areas underlain by the impure limestones, shales, and cherts of the Warsaw Limestone or Fort Payne Formation. The most extensive barrens areas are in the Eastern Highland Rim and Pennyroyal Plateau, but they occur throughout the Highland Rim (Wolfe, 1996a). In general, the Western Highland Rim is the most dissected of the three subdivisions, the Pennyroyal Plateau is the least dissected, and the Eastern Highland Rim is the most topographically variable (slope, table 4).
The Cumberland Plateau is a broad upland with nearly flat topography locally broken by deeply incised stream valleys and low hills. The general elevation of the plateau is about 600 m and is bounded by prominent escarpments that descend roughly 300 m to the Valley and Ridge Physiographic Province (Fenneman, 1938) and the Highland Rim, respectively east and west of the Plateau. The main surface of the Cumberland Plateau is capped with a sequence of Pennsylvanian sandstones, shales, and conglomerates. These siliciclastic rocks are generally unaffected by karst processes and are typically on the order of 100 m thick. The siliciclastic caprock is underlain by a sequence of Upper Mississippian rocks consisting of a transition sequence of shales, sandstones, and impure limestones (the Pennington Formation). The Pennington Formation, in turn, overlies thick (greater than 30 m), relatively pure limestones (the Bangor, Monteagle, and St. Louis Limestones) separated by thinner confining units of sandstone, shale, or impure limestones (Miller, 1974; White and White, 1983; Crawford, 1987; fig. 9). The Mississippian limestones crop out along the escarpments that bound the Cumberland Plateau and in deeply incised valleys, locally known as "coves," where stream erosion has cut through the caprock. By far the largest and most prominent stream valley incised into the plateau is that of the Sequatchie River (Milici, 1968; Miller, 1974; Crawford, 1989). Karst processes and features on the Cumberland Plateau are concentrated almost exclusively along the escarpments and coves.
The escarpments and coves along the edges of the Cumberland Plateau and along its transition to the Sequatchie Valley form one of the classic karst areas of North America (White and White, 1983; Crawford, 1987, 1989, 1992). The relatively thick, impervious cap of siliciclastic rocks concentrates recharge at points where the cap is breached by erosion. Under the influence of concentrated recharge, driven by several hundred meters of relief, the thick, relatively pure limestones have developed highly efficient conduit systems capable of accommodating perennially flowing streams. Along the face of the escarpments and the valley walls, the recurring sequence of cavernous limestone alternating with shaley or cherty confining units produces a characteristic pattern of ground-water/surface-water interaction (fig. 9). Numerous springs emerge where confining units crop out, then re-enter the ground-water system through dissolution openings in the next limestone unit down the sequence (White and White, 1983; Crawford, 1989). The combination of concentrated recharge, high relief, low aquifer storage, and well-developed conduit systems make the karst aquifers of the Cumberland Plateau vulnerable to contamination wherever the caprock has been breached.
The floor of the Sequatchie Valley is cut by the Sequatchie Valley (thrust) Fault (Milici, 1968). West of the fault, the rock units that crop out and the general physiographic relations are broadly analogous to those of the Eastern Highland Rim. East of the fault, the floor of the Sequatchie Valley can be considered an extension of the Valley and Ridge karst region discussed in the next section.
The Valley and Ridge Physiographic Province (Fenneman and Johnson, 1946) is a linear belt of folded and faulted sedimentary rocks that extends about 2,000 km, trending northeast-southwest from the St. Lawrence Valley to the Coastal Plain in Alabama (Fenneman, 1938). The Valley and Ridge takes its name from the characteristic topographic sequence of alternating ridges and stream valleys whose structurally controlled orientations are roughly parallel to the axis of the entire province (Fenneman, 1938). In Tennessee, the Valley and Ridge occupies roughly 21,000 km2, making it the second largest karst region in Tennessee, after the Highland Rim.
The major characteristic that distinguishes Valley and Ridge karst from karst in the Nashville Basin and Highland Rim or along the Cumberland Plateau is the strongly folded and faulted structure (fig. 10). Large-scale structures, such as major folds and thrust faults, control the spatial arrangement of the various rock units and determine (1) which units crop out at the surface, (2) where weathering and ground-water recharge occur, and (3) where secondary porosity can best develop in rocks of suitable lithology (Hollyday and Hileman, in press). Dissolution openings and active karst development may extend to depths of 180 m or deeper in the Valley and Ridge, compared to 30 m in the outer Central Basin, because karst aquifers and confining units are not horizontal (Hollyday and Hileman, in press; Sid Jones, Tennessee Department of Environment and Conservation, oral commun., 1997).
Throughout much of the Valley and Ridge, thick accumulations of regolith store and transmit large quantities of ground water (Bailey and Lee, 1991; Hollyday and Hileman, in press). In some areas, the dominant direction of ground-water movement in regolith tends to be normal to strike and toward the structurally controlled valley bottoms, with valley streams acting as discharge points (Bailey and Lee, 1991). Elsewhere, relict bedding structures may cause water to flow along strike in the regolith (Sid Jones, Tennessee Department of Environment and Conservation, oral commun., 1997). Potential for ground water to move relatively long distances parallel to strike (along the valley axis) is greatest in carbonate units with well-developed conduit systems (Bailey and Lee, 1991).
Hollyday and Hileman (in press) divided the Valley and Ridge Physiographic Province into hydrogeologic terranes based on lithology and well yields. They noted three major karst terranes: limestone, dolomite, and argillaceous carbonate rock. Limestone, such as the Jonesboro Limestone, and dolomite, such as the Knox Dolomite, develop significant conduit porosity in the Valley and Ridge. Limestone forms broad valley bottoms with varying thicknesses of alluvial cover. Dolomite also crops out in valley bottoms, but some dolomite, notably the Knox, also forms knobby ridges with chert-capped residual cover tens of meters thick (E.F. Hollyday, U.S. Geological Survey, oral commun., 1996). Argillaceous carbonate rock, such as the Moccasin or Lenoir Limestones of the Chickamauga Group, contains sufficient carbonate minerals to develop dissolution-enlarged openings. However, the high concentrations of insoluble materials in these units limit conduit development and tend to clog dissolution-enlarged openings with insoluble residue (Hollyday and Hileman, in press).
The "western toe of the Blue Ridge" (Hinkle and Sterrett, 1976, 1978; Hollyday and Hileman, in press) is a special localized setting like the coves and escarpments of the Cumberland Plateau. The western toe occurs along the boundary between the Valley and Ridge and Blue Ridge Physiographic Provinces with isolated outliers scattered within the Blue Ridge. Although these areas are highly localized, they provide important sources of ground water (Leonard, 1962; Hinkle and Sterrett, 1976, 1978; Meng and others, 1985, p. 431; Hollyday and others, 1997; Hollyday and Hileman, in press) and have ecological significance as habitat for rare plants and animals (Killebrew and Safford, 1874; Tryon, 1992; Wolfe, 1994). The main features of the western toe settings are (1) topographic juxtaposition of a dolomite or (less commonly) limestone valley bottom with quartzite slopes; and (2) a sequence, from top to bottom, of coarse clastic alluvium and colluvium, fine-grained residuum, and fractured, cavernous carbonate rock (fig. 11).
The heavily folded and faulted structure of the Valley and Ridge has significant implications for the fate and transport of pollutants. Steeply dipping, dissolution-enlarged joints and bedding planes are potential pathways for penetration of DNAPL deep into the ground-water system. The lithologic conditions under which such deep migration of DNAPL is most likely to occur are the same conditions that have the highest potential for long-distance movement of ground water, and dissolved contaminants, parallel to strike.
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