Groundwater Movement

The flow of water through the subsurface, known as groundwater movement, obeys well-established principles that allow hydrologists to predict flow directions and rates.

Porosity

The movement of water through the Earth’s subsurface is only one part of a larger circulation system known as the hydrologic cycle. This cycle involves the continuous transfer of water between natural reservoirs within the physical environment, such as the oceans, polar ice caps, groundwater, surface water, and the atmosphere. The main processes in the hydrologic cycle are precipitation, evaporation, transpiration by plants, surface-water runoff, and subsurface groundwater flow. When precipitation falls to the land surface as rain, snow, or other form, some of this water runs off into streams, some evaporates back into the atmosphere, and the remainder soaks into the ground. The water that infiltrates the land surface is taken up and transpired by plants, or percolates deeper to eventually become groundwater. For water to percolate into the subsurface, interconnected void spaces must be available between particles in the underlying geologic materials. The ratio of void space to the total rock or soil volume is called porosity. This proportion is usually expressed as a percentage. The higher this ratio, the more void space there is to hold water.

There are various types of porosity. Unconsolidated materials such as soil and sediment have pore spaces between adjacent grains, referred to as intergranular porosity. The ratio of pore space to total volume depends on several factors, including particle shape, sorting, and packing. Loosely packed sediments composed of well-sorted, spherical grains are the most porous. Porosity decreases as the angularity of the grains increases because the particles pack more closely together and the protuberances on the particles due to their irregular shape fill in more of the void space. Similarly, as the degree of sorting decreases, the pore spaces between larger grains become filled with smaller grains, and porosity decreases. Values of porosity for unconsolidated materials range from 10 percent for unsorted mixtures of sand, silt, and gravel to about 60 percent for some clay deposits. Typical porosity values for uniform sands are between 30 percent and 40 percent.

Rocks have two main types of porositypore spaces between adjacent mineral grains and voids that are a result of fractures. Rocks formed from sedimentary deposits (such as shale and sandstone) may have significant intergranular porosity, but it is usually less than the porosity of the sediments from which they were derived. This dichotomy is a result of the compaction and cementation that takes place during the process of transforming sediments into rock. Therefore, although sandstone porosities may be as high as 40 percent, they are commonly closer to 20 percent because of the presence of natural mineral cements that partially fill available pore spaces. Igneous and metamorphic rocks are composed of tightly interlocked mineral grains and, therefore, have little intergranular porosity. Virtually all void space in such rocks is a result of fractures (joints and faults). For example, granite (a dense, igneous rock) usually has a porosity of less than 1 percent, but porosity may reach 10 percent if the rock is fractured.

There are additional types of porosity that occur only in certain kinds of rocks. Limestone, a rock that is soluble in water, especially if the water is acidic, can develop solution conduits, or channels, along fractures and bedding planes. Given enough time, solution weathering may lead to the development of a cave, which has 100 percent porosity. The overall porosity of solution-weathered limestone sometimes reaches 50 percent. Rocks created by volcanic eruptions may contain void space in the forms of vesicles (cavities left by gases escaping from lava), vertical shrinkage cracks developed during cooling (known as columnar joints), and tunnels created by flowing lava (called lava tubes). In extreme cases, the porosity of volcanic rocks may exceed 80 percent.

Underground water includes all water that exists below the land surface, but the subject of groundwater movement is mainly concerned with the water that occurs in the zone of saturation, where all empty spaces are completely filled with water. Between the zone of saturation and the land surface, void spaces contain mostly air, unless a heavy rainfall or a period of snowmelt has just occurred. Water in this upper zone is held under tension by attractive forces between soil particles and water molecules (surface tension forces and molecular polarity effects).

The water table forms the uppermost surface of the zone of saturation and is characterized by having a water pressure equal to atmospheric pressure. It ranges in depth from very close to the land surface in humid regions to hundreds of meters below the surface in desert environments. In general, the water table mimics the surface topography but with more subdued slopes. When the water table intersects with the surface topography, a surface water feature such as a lake, swamp, river, or spring results. Below the water table, in the zone of saturation, geologic materials are completely saturated, and water pressure (called hydrostatic pressure) increases with depth in the same manner that pressure increases with depth in an open body of water. Water contained within the zone of saturation is generally called groundwater. It is this zone that supplies water to a well when it occurs in a particular type of geologic formation known as an aquifer.

Groundwater Energy

To understand groundwater flow, it is necessary to examine the forms of energy contained in groundwater. The total energy in any water mass consists of three components: elevation head, pressure head, and velocity head. The elevation head represents the potential energy of the water due to its elevation above mean sea level. The pressure head represents the potential energy of the water due to the hydrostatic pressure of the surrounding fluids. The velocity head corresponds to the kinetic energy of the water resulting from its physical movement. Because groundwater moves relatively slowly, velocity head can usually be neglected, leaving the total energy essentially equal to the sum of the elevation head and the pressure head. This quantity is known as the hydraulic head. Thus the hydraulic head of a given water particle varies directly with its elevation (usually expressed in meters above mean sea level) and its hydrostatic pressure.

The water table can be considered a surface with a variable hydraulic head. Because water pressure at the water table is always known, since it is by definition equal to atmospheric pressure, the change in hydraulic head across the water table depends only upon the variation in elevation. Below the water table, the hydraulic head depends on both the elevation and the water pressure, which increases with depth. Therefore, the variation in the hydraulic head with depth below the water table reflects the relationship between decreasing elevation head and increasing pressure head. If the increase in hydrostatic pressure exactly offsets the decrease in elevation head, then the hydraulic head will not change with depth.

Groundwater moves in response to differences in hydraulic head between two locations. The direction of movement is always from areas of higher hydraulic head toward areas of lower hydraulic head. The change in hydraulic head over a specified distance is known as the hydraulic gradient. Both horizontal and vertical hydraulic gradients can exist.

Horizontal hydraulic gradients are usually defined by the change in water table elevation between any two locations. As water moves through the subsurface, it flows along the steepest hydraulic gradient. Therefore, it is possible to determine the compass direction of groundwater movement from a knowledge of how the water table elevation varies over distance. Because the water table often mimics the land surface, general groundwater flow directions can sometimes be predicted on the basis of surface topography.

Vertical hydraulic gradients describe changes in hydraulic head with depth. As groundwater flows in any given horizontal direction, it may also be rising or sinking, depending on the vertical hydraulic gradient. In locations where hydraulic head decreases with depth below the water table, groundwater flow has a downward component, resulting in recharge areas. Where hydraulic head increases with depth, groundwater flow has an upward component, creating a discharge area. Recharge areas commonly occur in the higher elevations of a particular landscape, and discharge areas usually occur in the valleys near lakes, streams, and swamps. This year-round flow of groundwater from higher to lower elevations permits streams to flow in the dry summer months when there is little surface runoff. In certain situations, the water pressure conditions can cause groundwater to move “uphill” with respect to the surface topography. Therefore, the land surface is not always a good indicator of groundwater flow directions.

Groundwater movement can be divided into local and regional flow. In areas of rugged topography, most groundwater flow is local, meaning that it moves from the hilltops to the nearest stream or lake. In more gentle terrains, however, or in areas where the zone of groundwater movement is very thick, some flow escapes the local system into a deeper, regional system. Thus, water may enter the subsurface at a local zone of recharge and move long distances before surfacing at a regional discharge area. Identifying the boundaries of local and regional flow systems requires detailed information about the horizontal and vertical distributions of hydraulic head over a large area.

Rate of Flow and Permeability

The rate at which groundwater moves through the subsurface can also be determined on the basis of scientific principles. Groundwater flow velocities depend on two factors: the hydraulic gradient and the permeability of the geologic materials involved. Considering that differences in hydraulic head are the driving force behind groundwater movement, it is a logical consequence that the hydraulic gradient is related to the rate of flow. All other things being equal, the steeper the gradient, the faster groundwater will move. Hydraulic gradients may change throughout the year, reflecting the influences of recharge rates. In the spring, when recharge is high, the water table will rise fastest beneath recharge areas, producing a steeper hydraulic gradient and higher flow velocities. As the water table drops throughout the summer, the hydraulic gradient also declines, leading to slower groundwater flow.

Permeability is defined as the ability of porous formations to transmit fluids, and is a property of the geologic material in question. This property depends on both the size of void spaces and the degree to which they are interconnected. Thus, some high-porosity materials such as clay (up to 60 percent porosity) and pumice (a vesicular volcanic rock with up to 87 percent porosity) have very low permeability because the void spaces within them are largely isolated from one another. Materials that have high permeability include sand, gravel, sandstone, and solution-weathered limestone. Rocks with low porosities, such as shale, quartzite, granite, and other dense, crystalline rocks, typically also have low permeabilities, unless they are significantly fractured.

Groundwater is forced to move along tortuous paths through geologic materials as it follows the connecting spaces between voids. Therefore, even in highly permeable materials, groundwater flows much more slowly than does surface water in a river. Whereas the velocity of stream flow may be measured in meters per second, groundwater velocities commonly range from 1 meter per day to less than 1 meter per year (averaging about 17 meters per year in rocks). The highest velocities occur in rocks that are heavily fractured. In the extreme case, groundwater can actually move as a subsurface stream through cavernous limestone or volcanic rock. This situation is not normal, however, and underground streams are much less common than the average person might suspect.

Study of Groundwater Movement

The first step in studying groundwater movement in a particular area is to determine the hydraulic gradients that exist. Horizontal gradients are defined by measuring groundwater elevations in wells that intersect the water table. To define vertical gradients that may be present, piezometers are used. Piezometers are essentially vertical pipes that are open at the bottom to allow water to enter. These devices enable the hydraulic head at a particular depth below the water table to be determined. The elevation head is the elevation at the bottom of the piezometer, and the pressure head is the height to which water rises in the piezometer above the intake point.

Contour maps of the variation in hydraulic head are prepared from data collected from numerous wells and piezometers. These maps can show either horizontal variations (such as a map of the water table) or vertical changes in hydraulic head (as shown in a cross-sectional view). The contour lines used on these maps to connect points of equal hydraulic head are called equipotential lines. Groundwater flow directions can be determined once the distribution of known hydraulic head values has been contoured. Flow lines, depicting the idealized paths taken by water particles, are drawn to intersect equipotential lines at right angles, indicating that groundwater moves along the steepest hydraulic gradient. The resulting gridlike pattern of equipotential and flow lines (called a flow net) is a two-dimensional representation of the groundwater flow system.

Flow nets are much better indicators of groundwater flow directions than the land surface topography. Also, from a study of flow nets, it may be possible to determine accurately recharge and discharge areas. In the map view, flow lines will diverge from areas of recharge and converge in discharge areas. In the cross-sectional view, flow lines will have downward components where recharge is occurring and upward components in discharge areas. If hydraulic head does not change with depth, flow lines will be horizontal, indicating that neither recharge nor discharge conditions exist. Flow nets do not give estimations of flow velocities unless permeability values are known for the materials involved. If an estimation of permeability can be obtained, groundwater flow velocity may be calculated by multiplying the permeability value and the hydraulic gradient and dividing this product by the porosity of the geologic formation. This calculation yields an average linear velocity of groundwater flow through the open area provided by void space.

Groundwater flow directions and velocities can also be determined by introducing a “tracer” into the groundwater and monitoring its migration through observation wells. This technique is especially helpful in areas of fractured rock, in which flow patterns are difficult to predict. Tracers are easily detectable substances that will dissolve readily and move with the groundwater without reacting with the geologic materials. Ideally, tracers are safe to use, inexpensive, and easy to detect in low concentrations. Examples of tracers used in groundwater studies include salts (sodium or potassium chloride), fluorescent dyes, and the radioactive isotopes of certain elements (helium, hydrogen, and iodine). The choice of tracer usually depends on the subsurface materials that it must pass through.

In areas where the hydrogeologic conditions are well defined through wells, piezometers, borings, and permeability tests, the groundwater flow system can be studied using computer models. Computers are used to calculate hydraulic head values for the area modeled given the rates of recharge (infiltration) and discharge (by evapotranspiration or discharge to a stream, lake, or well). These computed head values can then be contoured to create a flow net. Although computers are powerful tools, the accuracy of their predictions cannot surpass the accuracy of the information provided for the model.

Significance

The understanding of groundwater movement is important to the utilization and conservation of aquifers for water supply. From a quantity point of view, flow rates determine which geologic materials will serve as suitable groundwater sources. The rates of flow must equal or exceed the desired pumping rate for a well to be successful. In this regard, permeability is the limiting factor, because the drawdown (depletion of water) from a well due to pumping will always create a sufficient hydraulic gradient to favor flow toward the well. If the permeability is too low to support the pumping rate, the geologic formation will become dewatered. For domestic wells requiring only 7 to 19 liters per minute, even moderately impermeable materials may supply sufficient water to be considered an aquifer. High-yield wells (exceeding 190 liters per minute), however, can be sustained only in very permeable materials, such as sand and gravel, sandstone, and solution-weathered limestone.

Understanding groundwater flow is also important to water quality. As groundwaters move through the subsurface, they dissolve minerals from the geologic materials with which they have contact. Therefore, groundwater that has traveled great distances in a regional flow system will tend to be the most mineralized and may be the least desirable for a drinking water source. With an increasing number of possible contamination sources, such as landfills, corrupted underground tanks, and accidental spills, the protection of groundwater supplies is crucial. A knowledge of groundwater flow directions and rates aids in predicting contaminant migration. It is particularly important to identify areas that recharge regional flow systems because these zones have the greatest potential impact if they become polluted. Municipalities concerned with preserving their well fields should undertake wellhead protection studies. The purpose of such studies is to delineate the well field recharge areas that need to be protected from any type of land use that could lead to groundwater contamination. Contamination of a drinking water source can have serious effects. An incident in Walkerton, Ontario, Canada, in 2000, in which surface runoff contaminated municipal water wells with the E. coli bacteria, resulted in the deaths of seven people and serious illness in another 2,300 individuals. As scientists discover the implications of groundwater contamination from industries, chemicals, and plastics, studying groundwater movement is increasingly important.

Contaminant migration is difficult to predict in fractured rock formations, especially in solution-weathered limestone. Because groundwater movement follows a sometimes random network of discontinuous openings in these settings, water levels may not be related from one well to the next. Therefore, maps of the water table for such areas are often impossible to construct. The flow of water through fractured rock is one of the concerns surrounding the choice of a repository for high-level nuclear wastes, which need to be isolated from the environment for at least ten thousand years.

Principal Terms

elevation head: the elevation of a given water particle above a certain point, usually mean sea level

equipotential line: a contour line connecting points of equal hydraulic head

groundwater: water found in the zone of saturation

hydraulic head: the sum of the elevation head and the pressure head at any given point in the subsurface

hydrostatic pressure: the pressure at any given point in a body of water at rest from the weight of the overlying water column

permeability: the ability of rock, soil, or sediment to transmit a fluid (commonly water)

porosity: the ratio, usually expressed as a percentage, of the total volume of void (empty) space in a given geologic material to the total volume of that material

pressure head: the height of a column of water that can be supported by the hydrostatic pressure at any given point in the subsurface

vadose zone: the region of soil between the surface and the water table in which void spaces contain both air and water

velocity head: the height to which the kinetic energy of fluid motion is capable of lifting that fluid

water table: the upper surface of the zone of saturation

zone of saturation: a subsurface zone in which all void spaces are filled with water

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