Overview and Nomenclature

Aquifers are geologic formations in the subsurface that can store & transmit water (Figures 1 and 2). As we will see later in the section titles Basic Aquifer Properties, there are specific rock and soil properties that govern these two functions. Adequate storage requires that there be sufficient void space between particles, in fractures, or generated by compressing the aquifer under pressure, to provide usable quantities of water. Adequate transmission requires that the void spaces where water occurs be well enough connected that it can percolate or flow under either natural or pumping-driven conditions, at a rate that will support sustained use.

Figure 1. Photo of the Casper aquifer in Eastern Wyoming, composed of cemented medium sandstone. The aquifer serves as the sole water supply for the town of Laramie, WY, population ~27,000.

Source: Demian Saffer

Figure 2. Photo of a fractured limestone member of the Casper formation. Sedimentary layers, or beds, are near-horizontal, and primary fractures are vertical.

Source: Demian Saffer

These definitions are intentionally vague because they depend on the scale of intended use. For example, an aquifer that provides water for a large city will need to sustain higher pumping rates at wells (on order of tens of thousands of gallons per minute) than one that provides for a single-family (a few to perhaps ten gallons per minute). For reference, Penn State University relies almost exclusively on groundwater pumping for its water supply at the University Park campus, with a total extraction of ~2.5 million gallons per day (about 1750 gallons per minute) distributed among several pumping well fields.

In contrast to an aquifer, an aquitard, often also termed a confining layer or aquiclude, is a geologic formation in the subsurface that does not transmit water effectively – and therefore acts as a barrier to groundwater flow. In general, aquifers are usually composed of sediments or sedimentary rocks with grain sizes larger than fine- to medium sand (>~125 µm diameter), or of fractured rock. Aquitards are typically composed of fine-grained sediments or sedimentary rocks or those in which the pore spaces have been filled by mineral cements (silts, siltstones, shales, clays, cemented sandstones, or unfractured limestones).

Aquifer Anatomy

In the simplest sense, you might imagine an aquifer formation that may be covered by a veneer of soil, and which extends downward from a few feet below the surface for several tens or even hundreds of feet (Figure 3). At some depth below the land surface, the interstices between soil or sediment particles, or the fractures in the rock, will be water-filled or saturated. Shallower than that depth, these interstices, or pore spaces, will be filled with air, water vapor, and some liquid water bound to the surfaces of the rock (Figures 3-5). This zone is the unsaturated zone, also known as the “zone of soil moisture” or the vadose zone. The water table marks the top of the ground water system and is formally defined as the depth at which the pressure in the subsurface is equal to the atmospheric pressure. Immediately above the water table, there is a narrow zone of saturation termed the capillary fringe. In this zone, water is wicked upward in pore spaces due to capillary forces. This is analogous to capillary tube experiments you may have seen or performed in physics or chemistry classes in high school; it occurs due to interaction between the polar water molecule and the surfaces of the solids, and is directly related to the fact that water has a surface tension (as you may remember from Module 1!). In the capillary fringe, pores are saturated, but pressures are sub-atmospheric, meaning that the water is under suction as it is pulled or wicked upward. The nomenclature “fringe” reflects the fact that slight variations in grain size lead to variations in the height that water is drawn (again, think to the capillary tube experiment, and effects of different tube sizes) (Figure 5).

Any precipitation or surface water that infiltrates to the water table must percolate through the vadose zone in order to recharge the aquifer. As we will see later in this module infiltration and recharge typically constitute only a small fraction – rarely more than 10% - of precipitation, because most water that falls in events is returned to the atmosphere by transpiration or evaporation, becomes runoff (i.e., if the capacity for infiltration is exceeded by the rate of precipitation), or is bound by soils in the vadose zone.

Figure 3. Schematic diagram showing a vertical section from the land surface through the vadose zone, capillary fringe, and into an unconfined aquifer.

Source: From W.M. Alley et al., 1999, USGS Circular 1186: Sustainability of Ground Water Resources [1]

Figure 4. Illustration of capillary action; water is drawn to different heights in tubes of different internal diameters.

Source: From USGS Water Science School [2]. Photo Credit: C. Robinson, West Texas A&M Univ.

Figure 5. Cross-sectional diagram of a water Illustration of capillary action; water is drawn to different heights in tubes of different internal diameters, showing the vadose zone (unsaturated zone), water table, and saturated aquifer. Insets below show water occurrence in crevices or fractures in rock (left) or between sediment or soil grains (right).

Click Here for Text alternative of Cross-sectional diagram of water

Water (not ground water) held by molecular attraction surrounds surfaces of rock particles. All openings below the water table are full of ground water. The Unsaturated Zone is on the land surface and above the water table which contains the saturated zone, surface water, and ground water.

Source: U.S. Geological Survey, Water Science for Schools: Aquifers [3].

Porosity (usually denoted by the symbol η, which is Greek letter 'eta') is the primary aquifer property that controls water storage, and is defined as the volume of void space (i.e., that can hold water in the zone of saturation) as a proportion of the total volume (Figure 10).

Figure 10. Schematic diagrams illustrating porosity in sedimentary rock with two different particle, or grain sizes (left), and in a crystalline rock (right). In the latter example, the porosity is restricted to the boundaries between “grains” that remains after they crystallized or grew, and is usually <2-5%.

Source: Earthsci.org (Creative Commons Attribution-NonCommercial-ShareAlike 3.0 Unported License), after lecture notes [6] from S.A. Nelson.

Porosity is expressed as either a fraction, or a percentage:

$\eta =V_{voids}/V_{total}$ , or if reported as a percentage,

$$\eta =(V_{voids}/V_{total})x100$$

For aquifers composed of sedimentary rocks or sediments, porosity is usually in the range of ~10-35%. For unfractured crystalline rock, porosity is quite a bit lower - on order of a few percent - because there is little porosity between individual grains other than the vary narrow interfaces along their boundaries (Figure 10).

Several factors can affect porosity. In sedimentary rock and sediments, controls on porosity include sorting, cementation, overburden stress (related to burial depth), and grain shape. Poorly sorted sedimentary deposits, in which there is a wide distribution of grain sizes, typically have lower porosity than well-sorted ones (Figure 11). This is because the finer particles are able to fill in spaces between the larger grains. Cementation caused by precipitation of minerals (typically calcium carbonate or silica) at grain boundaries also reduces porosity (Figure 11). Angular grains generally allow more efficient packing of particles than rounded or spherical ones, also leading to slightly lower porosity. Finally, the more deeply sediments or sedimentary rock are buried, the larger the weight of the overburden; the higher stress leads to compaction, tighter packing of the grains, and lower overall porosity.

Secondary processes that act on the rock or sediment after its formation, primarily weathering by physical or chemical mechanisms, can also affect porosity. Physical weathering by wind or water movement can remove fine clay-sized particles from the sediment (a process termed winnowing), leading to increased porosity near the Earth’s surface. Chemical weathering of certain rock types can lead to clay and oxide formation; depending on the environment and initial composition of the aquifer grains, the clays and oxides may subsequently be removed (porosity increase), or they may grow at the boundaries of other particles and reduce the porosity.

In fractured rock (whether fractured crystalline or cemented sedimentary rock), porosity is typically ~2-5%. The pore space is almost entirely composed of the fractures or cracks themselves, which are typically a millimeter or less in aperture (Figure 12). Two primary factors control porosity – and the connectedness of porosity – in fractured rock. First, increased stress, related to the depth of burial and the weight of the overburden, exerts a clamping force that causes the closure of cracks or fractures (decreases fracture aperture). In some types of rock – most notably limestones – chemical weathering occurs via dissolution as water flows along and through fractures. This leads to increased fracture aperture. Significant enlargement of fractures can lead to the development of karst, typified by large open fractures, caves, and caverns, as well as sinkholes and hummocky topography that ensue as the underlying rock is gradually dissolved (Figure 13).

Interestingly, grain size does not affect porosity. For example, consider a box filled with spherical particles packed as tightly as possible. The proportion of empty space (porosity) would be the same whether the particles are marble-sized, pea-sized, or golf ball-sized; the porosity is controlled entirely by the geometry of the particles – not their dimensions. As we’ll see in the next section, however, grain size does strongly affect the ability of aquifers to transmit water because it directly controls the size of the pore spaces where the water percolates. For example, unconsolidated clays (grain sizes of a few to tens of microns) commonly have porosities of over 50-60%, but they transmit water only one-thousandth to one millionth as well as sands with porosities of 20-30% (grain sizes of a few hundred microns).

Specific Yield (denoted as Sy) is another important quantity for water storage in unconfined aquifers. Sy is defined as the proportion of water occupying void spaces that drains under gravity. Because some water is bound, or adsorbed, to the aquifer particles or fractures, the specific yield is always lower than the porosity (take a look at Figure 8 [7], inset at top left). The attraction between water molecules and the aquifer is due – you guessed it! – to the polar nature of water and surface tension. In sands and fractured rock, Sy is typically a large fraction (>90%) of the porosity, whereas in fine-grained sedimentary deposits Sy may be as low as a few percent because the surface area interacting with water molecules is higher, and pores are smaller, allowing the aquifer to retain more water. In unconfined aquifers, Sy controls the amount of water that can be extracted by pumping.

In confined aquifers, the compressibility of the aquifer is the dominant control on water storage and release. As described above (see Figure 8, right panel), when water is extracted from confined aquifers by pumping or flow to natural springs, the aquifer remains saturated, but the water pressure decreases. Upon depressurization, the aquifer itself can compress slightly. If water is recharged or injected, the opposite occurs: pressure increases and the aquifer expands very slightly. Essentially, by increasing water pressure, more water mass is being “crammed” into the pore space in the aquifer, and vice versa. Although exaggerated, one way to visualize this is to think of pores in the rock or aquifer as a juice box. By changing the pressure inside the box, it will expand or contract. In the same way that a soft juice box will deform more than a stiff one for a given change in pressure, a more compressible aquifer will yield more water than a stiffer aquifer, for the same depressurization. The storage of water in confined aquifers is termed the specific storage, and reported in units of Volume of water/Volume of aquifer per change in water level (so the units are 1/length; e.g., 1/m or 1/ft).

Figure 11. (Left) Poorly sorted sediment, with fine particles filling spaces between larger grains. (Right) Cementation in pore spaces (gray) that reduces the porosity.

Source: Earthsci.org (Creative Commons Attribution-NonCommercial-ShareAlike 3.0 Unported License), after lecture notes [6] from S.A. Nelson.

Figure 12. Schematic diagram illustrating the difference between intergranular porosity in sedimentary rock (top) and fracture porosity (bottom).

Click Here for Text alternative of Figure 12

Sedimentary aquifers with intergranular porosity have groundwater storage between grains of sediment. Sedimentary aquifers with fracture porosity have groundwater storage in solution cavities.

Source: Tasmanian Government - Dual Porosity Aquifers [8]

Figure 13. Schematic of a karst system.

Source: United States Geological Survey "Exploring Caves" teachers packet

Transmission

The ability of an aquifer to transmit water – or of an aquitard to slow the flow of water – is the second essential ingredient controlling groundwater movement. It is also the most variable in natural materials; distances in astronomy are the only other quantity in nature that varies over a similar range! For example, the difference in groundwater flow rate for shale vs. gravel is a factor of 1,000,000,000,000 (yup…one trillion). That’s the difference between the size of an iPhone and the distance from the Earth to the Sun.

Groundwater transport properties are described by two related quantities. Hydraulic Conductivity, denoted by K, is a measure of the ability of a particular fluid (usually water) to flow through the rock or sediment. Permeability, denoted by a lower-case k, is often also termed “intrinsic permeability” and describes the ability of the geologic formation alone to transmit fluid. Although related, the key difference is that hydraulic conductivity combines properties of the geologic formation and the fluid, whereas permeability describes only the rock properties. As described in Sidebar 1, the basic concept of hydraulic conductivity emerged from a series of ingenious experiments conducted in the mid-1880s by Henri Darcy, a French Engineer. These experiments led to Darcy’s Law, which forms the foundation for much of modern hydrogeology and petroleum engineering.

To illustrate the difference between K and k, consider the sandstone in Figure 14 below. The sandstone itself has a permeability, which is controlled by the size of the grains and pore spaces through which water can percolate, and the connectedness and geometry of the pores (more on that in a moment!). That permeability is a characteristic of the sandstone, regardless of whatever fluid might be moving through it, the temperature, or anything else. But the flow rate of water through this sandstone will be different than for oil, or for air, or any other fluid. So the same sandstone also has a hydraulic conductivity specific to a given fluid of interest.

Figure 14. Photo of the Aquia Creek Sandstone.

Source: USGS Multimedia Gallery.

Viscosity and Density

More specifically, it is the viscosity and density of the fluid that matter. More viscous fluids will flow more slowly through the same rock than less viscous ones. This is important for comparing different fluids (say, oil vs. water – whether you are thinking about an oil reservoir or contamination of groundwater by a gasoline spill). It is also important in considering the effects of temperature, because water viscosity decreases with increasing temperature: it’s less than half as viscous at 90° than at 32° F. So even for the same aquifer, the hydraulic conductivity goes up if it is warmer! This makes some sense – if the water is less viscous (i.e. “thinner”), it will flow more easily through the aquifer.

So…that’s how we define permeability and hydraulic conductivity. But what controls their magnitude? The main factors are grain size and shape, sorting, porosity (degree of compaction or fracture aperture), particle orientation or alignment that affects the tortuosity of the flow path, and cementation. Tortuosity is a measure of how far fluid must go to “circumnavigate” its way around particles: higher tortuosity indicates that water must go farther to get to its destination (a more tortuous path). For all of these mechanisms, the key underlying control on groundwater movement is the viscous resistance resulting from the interaction of the fluid with solid surfaces in the aquifer (grain edges or fracture walls).

Importance of Fractures

Fractured aquifers are one important and widely used class of aquifer because they are commonly both highly permeable and rapidly recharged. For example, groundwater recharge to the limestone aquifer beneath Nittany Valley in the Spring Creek watershed is around 30-45% of the annual precipitation (in comparison to typical recharge of <10% of precipitation). Fractured aquifers are permeable despite their overall low porosity (usually <5%) because natural fractures usually form in consistent orientations and are well connected in networks over hundreds of meters to tens of kilometers or more (Figures 15-16). The preferred orientation of major fractures leads to anisotropy in permeability, in which the aquifer may be more permeable parallel to the dominant fracture directions than in other orientations.

The rapid flow rates and direct pathways for recharge from the land surface also lead to concerns specific to fractured aquifers. In the absence of confining layers or thick soils, rapid recharge along fractures that extend to or nearly to the land surface increase vulnerability of contamination by surface activity, including fertilization of fields, pesticide application, or spills. Direct connections between surface water bodies and groundwater through major fracture systems also increase the potential for water-borne pathogens to enter the groundwater system, especially during periods of high flow or if confining layers along stream beds are breached. Compounding this risk, if contamination does occur, flow along fracture networks can be very rapid and the direction and rates of contaminant transport difficult to predict - unless the fracture network in the subsurface is extraordinarily well known, which is rarely the case. Because of their potential for contamination, fractured aquifers are a subject of highly active research, including dedicated large-scale field programs (e.g., check out the U.S. Geological Survey’s Mirror Lake project [9]).

Figure 15. The Lockatong mudstone formation in NJ, characterized by two sets of fractures perpendicular and parallel to bedding (layering). The Lockatong hosts a trichloroethylene plume (TCE) at the Naval Air Warfare Center Research Site in West Trenton.

Source: USGS Flow and Transport in Fractured Rock toxics program [10]

Figure 16. (Bottom) Outcrop of a fractured rock aquifer at Mirror Lake, NH.

Source: USGS Flow and Transport in Fractured Rock toxics program [10]

Valley and Ridge Aquifer System

The Valley and Ridge aquifer system extends SW-NE across Central PA, West VA, and VA (Figure 19, purple area), and is the main water supply for much of this region. It is composed of layered Paleozoic sedimentary strata (shales, sandstones, and limestones) that were folded and deformed by a series tectonic collisions over 200 million years ago. The modern valleys in the Valley and Ridge province have formed where limestone, which is most susceptible to erosion, was exposed in the core of anticlines, or upfolds (Figure 20). The more resistant sandstones and shales form the regional ridges, like Mt. Nittany and Bald Eagle Ridge.

The principal aquifer unit in this system is the fractured limestone that underlies the valleys. As noted above, because it is fractured, it recharges rapidly, has a high fracture permeability, and wells drilled along the fractures are highly productive (c.f. Figure 17). Recharge is focused on the flanks of the ridges, where runoff flows over the less permeable shale and sandstone units and enters the groundwater through fractures or sinkholes above the limestone at the valley edges. Groundwater flow is generally toward the center of the valleys, and springs commonly feed the surface water systems. The water is characterized by a high hardness (Mg and Ca content; we’ll cover this in more detail in Module 7), derived from limestone dissolution. Dissolution of the limestone has formed extensive karst features (caves, caverns, sinkholes) throughout the region.

Figure 19. Map showing the extent of the Valley and Ridge (purple), Atlantic Coastal Plain (green), Blue Ridge (orange), and Appalachian Plateau aquifer systems.

Source: USGS Groundwater Atlas

Figure 20. Block diagram showing the folded strata that form the Valley and Ridge aquifer system.

Click Here for Text Alternative to Figure 20

The Paleozoic rocks of the Appalachian Plateaus and the Valley and Ridge Provinces range from nearly flat-lying to intensely folded They are commonly separated from crystalline rocks of the Blue Ridge and the Piedmont Provinces by faults. The sediments of the Coastal Plain Province overlie older rocks and dip gently toward the ocean.

Source: USGS Water Atlas

Atlantic Coastal Plain Aquifer System

The Atlantic Coastal Plain aquifer system extends North-South along much of the Eastern portions of New Jersey, Delaware, Pennsylvania, Virginia, and North Carolina (Figure 19). It consists of a sequence of layered sedimentary aquifers (sands and gravels) separated by series of aquitards, all deposited starting around 100 million years ago and continuing today. The layers slope, or dip, to the East and extend offshore for tens of km beneath the continental shelf (Figure 21).

Figure 21. Layered sequence of major aquifer and aquitard units in the Atlantic Coastal Plain system. The sedimentary units dip and thicken to the East.

Source: USGS Water Atlas

Recharge occurs by both natural and managed infiltration on land across much of the coastal plain; groundwater flow in the subsurface is mainly to the East along the sediment layers. One interesting consequence of this flow pattern is that there may be a sizable freshwater resource offshore that could be accessed by drilling in relatively shallow water on the continental shelf. During the last ice age, when conditions were substantially wetter than today and a nearly mile-thick ice sheet covered the northern extent of the aquifer system, recharge was probably even larger - and thus may have forced fresh water several tens of km offshore, where that “fossil” water may remain today!

The Atlantic Coastal Plain system is an important water source for domestic/municipal supply and industry in population centers throughout coastal North Carolina, Maryland, Virginia, Delaware, and New Jersey. However, concentrated, localized pumping has led to a reversal of flow direction (toward the wells instead of Eastward) in some of the aquifer units throughout the region. In addition to overarching concerns about the sustainability of withdrawals that exceed recharge rates, the flow reversal has led to local salt-water intrusion, whereby saline ocean water infiltrates the aquifer and in some cases renders it non-potable.

Central Valley/San Joaquin Valley Aquifer System

The Central Valley Aquifer system of Central CA lies in a large structural basin running approximately North-South, between the Coast Ranges to the West and the Sierra Nevada mountains to the East (Figure 22). The deep elongate basin is infilled with marine and continental sediments, primarily composed of interlayered sands and clays. The basin itself is formed by tectonic processes caused by East-West extension (these are the same forces that are causing continued uplift of the major mountain chains throughout the Basin and Range province of the Southwestern US, and which are one major control on orographic precipitation patterns in that region).

The continental deposits (Figure 22, orange) comprise the main aquifer units and range from one-half to over two miles in thickness. As is the case in the Valley and Ridge, the recharge is primarily focused around the valley perimeter as runoff over the flanks of surrounding high topography infiltrates and enters the groundwater system. Groundwater flow is primarily inward, toward valley center, with a component of flow down-valley to the North, parallel to surface water flow in the San Joaquin River.

Figure 22. (Top) Block diagram of the Central Valley Aquifer System.

Source: U.S. Geological Society

Figure 23. (Left) Contour map showing the thickness of continental deposits that form the aquifer units in the system. (Right) Map showing a generalized pattern of groundwater recharge and discharge from the aquifer system.

Source: U.S. Geological Society

The thick sedimentary sequence has formed a vast expanse of flat topography on the natural floodplain of the San Joaquin River. This, in combination with a mild climate that allows a year-round growing season, has made the Central Valley one of the most productive and largest agricultural centers in the world. The Central Valley aquifer system is highly utilized, primarily to augment limited allocations of surface water for irrigation. Since the mid-1920s, groundwater withdrawals have generally outpaced natural recharge to the aquifer, leading to dropping water levels, irreversible aquifer compaction, and land subsidence (as will be discussed in more detail next week, in Module 6.2). Until recently, groundwater withdrawals were neither heavily monitored nor regulated. However, in the face of an ongoing multi-year drought, a 2014 bill was signed into law that restricts pumping and implements groundwater sustainability plans (See What to Know about California's New Groundwater Law [14]; see also New California Groundwater Pumping Rules Signed Into Law [15]). Shallow aquifer units in the valley are also plagued by a wide range of water quality concerns associated with irrigation and return flow of irrigation water to the aquifer via infiltration; these include leaching of selenium, boron, and other constituents from soils; salinization; and high concentrations of pesticides and fertilizers. We’ll discuss all of these issues in more detail in upcoming modules about water quality and the effects of climate change.