Groundwater Problems

Groundwater Problems

Since prehistoric times, groundwater has been an important resource that people have relied on for drinking, irrigation, and industry. Groundwater feeds the lushness of desert oases in the Sahara, the amber grain in the North American high plains, and the growing cities of sunny arid regions. 

Though groundwater accounts for about 95% of the liquid freshwater on the planet, accessible groundwater cannot be replenished quickly, and this leads to shortages. Groundwater contamination is also a growing tragedy. Such pollution, caused when toxic wastes and other impurities infiltrate down to the water table, may be invisible to us but may ruin a water supply for generations to come. In this section, we’ll take a look at problems associated with the use of groundwater supplies. 

Depletion of Groundwater Supplies 

Is groundwater a renewable resource? In a time frame of 10,000 years, the answer is yes, for the hydrologic cycle will eventually resupply depleted reserves. But in a time frame of 100 to 1,000 years the span of a human lifetime or a civilization groundwater in many regions may be a non-renewable resource. By pumping water out of the ground at a rate faster than nature replaces it, people are effectively “mining” the groundwater supply. In fact, in portions of the desert Sunbelt region of the United States, supplies of young groundwater have already been exhausted, and deep wells now extract 10,000-year old groundwater. Some of this ancient water has been in rock so long that it has become too mineralized to be usable. A number of other problems accompany the depletion of groundwater.

Effects of human modification of the water table.

• Lowering the water table: When we extract groundwater from wells at a rate faster than it can be resupplied by nature, the water table drops. First, a cone of depression forms locally around the well; then the water table gradually becomes lower in a broad region. As a consequence, existing wells, springs, and rivers, and swamps dry up (figure above a, b). To continue tapping into the water supply, we must drill progressively deeper. Notably, the water table can also drop when people divert surface water from the recharge area. Such a problem has developed in the Everglades of southern Florida, a huge swamp where, before the expansion of Miami and the development of agriculture, the water table lay at the ground surface (figure above c, d). Diversion of water from the Everglades’ recharge area into canals has significantly lowered the water table, causing parts of the Everglades to dry up.
• Reversing the flow direction of groundwater: The cone of depression that develops around a well creates a local slope to the water table. The resulting hydraulic gradient may be large enough to reverse the flow direction of nearby groundwater (figure below a, b). Such reversals can allow contaminants, seeping out of a septic tank, to contaminate the well.
• Saline intrusion: In coastal areas, fresh groundwater lies in a layer above saline (salty) water that entered the aquifer from the adjacent ocean (figure below c, d). Because fresh water is less dense than saline water, it floats above the saline water. If people pump water out of a well too quickly, the boundary between the saline water and the fresh groundwater rises. And if this boundary rises above the base of the well, then the well will start to yield useless saline water. Geologists refer to this phenomenon as saline intrusion. 
• Pore collapse and land subsidence: When groundwater fills the pore space of a rock or sediment, it holds the grains apart, for water cannot be compressed. The extraction of water from a pore eliminates the support holding the grains apart, because the air that replaces the water can be compressed. As a result, the grains pack more closely together. Such pore collapse permanently decreases the porosity and permeability of a rock, and thus lessens its value as an aquifer (figure below e, f).

Some causes of groundwater problems.

Pore collapse also decreases the volume of the aquifer, with the result that the ground above the aquifer sinks. Such land subsidence may cause fissures at the surface to develop and the ground to tilt. Buildings constructed over regions undergoing land subsidence may themselves tilt, or their foundations may crack. In the San Joaquin Valley of California, the land surface subsided by 9 m between 1925 and 1975, because water was removed to irrigate farm fields.

Natural Groundwater Quality 

Much of the world’s groundwater is crystal clear, and pure enough to drink right out of the ground. Rocks and sediment are natural filters capable of removing suspended solids these  solids get trapped in tiny pores or stick to the surfaces of  clay flakes. In fact, the commercial distribution of bottled groundwater (“spring water”) has  become a major business worldwide. But dissolved chemicals, and in some cases methane, may make some natural groundwater unusable. For example, groundwater that has passed through salt-containing strata may become salty and unsuitable for irrigation or drinking. Groundwater that has passed through limestone or dolomite contains dissolved calcium (Ca2 ) and magnesium (Mg2 ) ions; this water, called hard water, can be a problem because carbonate minerals precipitate from it to form “scale” that clogs pipes. Also, washing with hard water can be difficult because soap won’t develop a lather. Groundwater that has passed through iron-bearing rocks may contain dissolved iron oxide that precipitates to form rusty stains. Some groundwater contains dissolved hydrogen sulphide, which comes out of solution when the groundwater rises to the surface; hydrogen sulphide is a poisonous gas that has a rotten-egg smell. In recent years, concern has grown about arsenic, a highly toxic chemical that enters groundwater when arsenic-bearing minerals dissolve in groundwater. 

Human-Caused Groundwater Contamination 

Contamination plumes in groundwater.

As we’ve noted, some contaminants in groundwater occur naturally. But in recent decades, contaminants have increasingly been introduced into aquifers because of human activity (figure above a). These contaminants include agricultural waste (pesticides, fertilizers, and animal sewage), industrial waste (dangerous organic and inorganic chemicals), effluent from “sanitary” landfills and septic tanks (including bacteria and viruses), petroleum products and other chemicals that do not  dissolve in water, radioactive waste (from weapons manufacture, power plants, and hospitals), and acids leached from sulfide minerals in coal and metal mines. The cloud of contaminated groundwater that moves away from the source of contamination is called a contaminant plume (figure above b).

The best way to avoid such groundwater contamination is to prevent contaminants from entering groundwater in the first place. This can be done by placing contaminants in sealed containers or on impermeable bedrock so that they are isolated from aquifers. If such a site is not available, the storage area should be lined with plastic or with a thick layer of clay, for the clay not only acts as an aquitard, but it can bond to contaminants. Fortunately, in some cases, natural processes can clean up groundwater contamination. Chemicals may be absorbed by clay, oxygen in the water may oxidize the chemicals, and bacteria in the water may metabolize the chemicals, thereby turning them into harmless substances. 

Where contaminants do make it into an aquifer, environmental engineers drill test wells to determine which way and how fast the contaminant plume is flowing; once they know the flow path, they can close wells in the path to prevent consumption of contaminated water. Engineers may attempt to clean the groundwater by drilling a series of extraction wells to pump it out of the ground. If the contaminated water does not rise fast enough, engineers drill injection wells to force clean water or steam into the ground beneath the contaminant plume (figure above c). The injected fluids then push the contaminated water up into the extraction wells. 

More recently, environmental engineers have begun exploring techniques of bioremediation: injecting oxygen and nutrients into a contaminated aquifer to foster growth of bacteria that can react with and break down molecules of contaminants. Needless to say, cleaning techniques are expensive and generally only partially effective.

Unwanted Effects of Rising Water Tables 

We’ve seen the negative consequences of sinking water tables, but what happens when the water table rises? Is that necessarily good? Sometimes, but not always. If the water table rises above the level of a house’s basement, water seeps through the foundation and floods the basement floor. Catastrophic damage occurs when a rising water table weakens the base of a hillslope or a failure surface underground triggers landslides and slumps. 
Figures credited to Stephen Marshak.

Hot Springs and Geysers

Hot Springs and Geysers

 Geothermal waters and examples of their manifestation in the landscape.

Hot springs, springs that emit water ranging in temperature from about 30° to 104°C, are found in two geologic settings. First, they occur where very deep groundwater, heated in warm bedrock at depth, flows up to the ground surface. This water brings heat with it as it rises. Such hot springs form in places where faults or fractures provide a high-permeability conduit for deep water, or where the water emitted in a discharge region followed a trajectory that first carried it deep into the crust. Second, hot springs develop in geothermal regions, places where volcanism currently takes place or has occurred recently, so that magma and/or very hot rock resides close to the Earth’s surface (figure above a). Hot groundwater dissolves minerals from rock that it passes through because water becomes a more effective solvent when hot, so people use the water emitted at hot springs as relaxing mineral baths (figure above b). Natural pools of geothermal water may become brightly coloured the gaudy greens, blues, and oranges of these pools come from thermophyllic (heat-loving) bacteria and archaea that thrive in hot water and metabolize the sulphur containing minerals dissolved in the groundwater (figure above c). 

Numerous distinctive geologic features form in geothermal regions as a result of the eruption of hot water. In places where the hot water rises into soils rich in volcanic ash and clay, a viscous slurry forms and fills bubbling mud pots. Bubbles of steam rising through the slurry cause it to splatter about in goopy drops. Where geothermal waters spill out of natural springs and then cool, dissolved minerals in the water precipitate, forming colourful mounds or terraces of travertine and other chemical sedimentary rocks (figure above d).

Under special circumstances, geothermal water emerges from the ground in a geyser (from the Icelandic spring, Geysir, and the word for gush), a fountain of steam and hot water that erupts episodically from a vent in the ground (figure above e). To understand why a geyser erupts, we first need a picture of its underground plumbing. Beneath a geyser lies a network of irregular fractures in very hot rock; groundwater sinks and fills these fractures. Heat transfers from the rock to the groundwater and makes the water’s temperature rise. Since the boiling point of water (the temperature at which water vaporizes) increases with increasing pressure, hot groundwater at depth can remain in liquid form even if its temperature has become greater than the boiling point of water at the Earth’s surface. When such “superheated” groundwater begins to rise through a conduit toward the surface, pressure in it decreases until eventually some of the water transforms into steam. The resulting expansion causes water higher up to spill out of the conduit at the ground surface. When this spill happens, pressure in the conduit, from the weight of overlying water, suddenly decreases. A sudden drop in pressure causes the super-hot water at depth to turn into steam instantly, and this steam quickly rises, ejecting all the water and steam above it out of the conduit in a geyser eruption. Once the conduit empties, the eruption ceases, and the conduit fills once again with water that gradually heats up, starting the eruptive cycle all over again. 

Figures credited to Stephen Marshak.

Tapping Groundwater Supplies

Tapping Groundwater Supplies 

We can obtain groundwater at wells or springs. Wells are holes that people dig or drill to obtain water. Springs are natural outlets from which groundwater flows. Wells and springs provide welcome sources of water but must be treated with care if they are to last.

Wells 

Pumping groundwater at a normal well affects the water table.

In an ordinary well, the base of the well penetrates an aquifer below the water table (figure above a). Water from the pore space in the aquifer seeps into the well and fills it to the level of the water table. Drilling into an aquitard, or into rock that lies above the water table, will not supply water, and thus yields a dry well. Some ordinary wells are seasonal and function only during the rainy season, when the water table rises. During the dry season, the water table lies below the base of the well, so the well is dry.

To obtain water from an ordinary well, you either pull water up in a bucket or pump the water out. As long as the rate at which groundwater fills the well exceeds the rate at which water is removed, the level of the water table near the well remains about the same. However, if users pump water out of the well too fast, then the water table sinks down around the well, in a process called drawdown, so that the water table becomes a downward-pointing, cone-shaped surface called a cone of depression (figure above b, c). Drawdown by a deep well may cause shallower wells that have been drilled nearby to run dry. 

Artesian wells, where water rises from the aquifer without pumping.

An artesian well, named for the province of Artois in France, penetrates confined aquifers in which water is under enough pressure to rise on its own to a level above the surface of the aquifer. If this level lies below the ground surface, the well is a nonflowing artesian well. But if the level lies above the ground surface, the well is a flowing artesian well, and water actively fountains out of the ground (figure above a). Artesian wells occur in special situations where a confined aquifer lies beneath a sloping aquitard. 

We can understand why artesian wells exist if we look first at the configuration of a city water supply (figure above b). Water companies pump water into a high tank that has a significant hydraulic head relative to the surrounding areas. If the water were connected by a water main to a series of vertical pipes, pressure caused by the elevation of the water in the high tank would make the water rise in the pipes until it reached an imaginary surface, called a potentiometric surface, that lies above the ground. This pressure drives water through water mains to household water systems without requiring pumps. In an artesian system, water enters a tilted, confined aquifer that intersects the ground in the hills of a high-elevation recharge area (figure above c). The confined groundwater flows down to the adjacent plains, which lie at a lower elevation. The potentiometric surface to which the water would rise, were it not confined, lies above this aquifer. Pressure in the confined aquifer pushes water up a well.

Springs 

Many towns were founded next to springs, places where groundwater naturally flows or seeps onto the Earth’s surface, for springs can provide fresh, clear water for drinking or irrigation, without the expense of drilling or digging. Some springs spill water onto dry land. Others bubble up through the bed of a stream or lake. Springs form under a variety of conditions: 

Geological settings in which springs form.

• Where the ground surface intersects the water table in a discharge area (figure above a); such springs typically occur in valley floors, where they may add water to lakes or streams. 
• Where flowing groundwater collides with a steep, impermeable barrier, and pressure pushes it up to the ground along the barrier (figure above b). 
• Where a perched water table intersects the surface of a hill (figure above c).
• Where downward-percolating water runs into a relatively impermeable layer and migrates along the top surface of the layer to a hillslope (figure above d). 
• Where a network of interconnected fractures channels groundwater to the surface of a hill (figure above e). 
• Where the ground surface intersects a natural fracture (joint) that taps a confined aquifer in which the pressure is sufficient to drive the water to the surface; such an occurrence is an artesian spring. 

Springs can provide water in regions that would otherwise be uninhabitable. For example, oases in deserts may develop around a spring. An oasis is a wet area, where plants can grow, in an otherwise bone-dry region.
Figures credited to Stephen Marshak.

Groundwater Flow

Groundwater Flow

What happens to groundwater over time? Does it just sit, unmoving, like the water in a stagnant puddle, or does it flow and eventually find its way back to the surface? Countless measurements confirm that groundwater enjoys the latter fate groundwater indeed flows, and in some cases it moves great distances underground. Let’s examine factors that drive groundwater flow. 

In the unsaturated zone the region between the ground surface and the water table water percolates straight down, like the water passing through a drip coffee maker, for this water moves only in response to the downward pull of gravity. But in the zone of saturation the region below the water table water flow is more complex, for in addition to the downward pull of gravity, water responds to differences in pressure. Pressure can cause groundwater to flow sideways, or even upward. (If you've ever watched water spray from a fountain, you've seen pressure pushing water upward.) Thus, to understand the nature of groundwater flow, we must first understand the origin of pressure in groundwater. For simplicity, we’ll consider only the case of groundwater in an  unconfined aquifer. 

The shape of water table beneath hilly topography.

Pressure in groundwater at a specific point underground is caused by the weight of all the overlying water from that point up to the water table. (The weight of overlying rock does not contribute to the pressure exerted on groundwater, for the contact points between mineral grains bear the rock’s weight.) Thus, a point at a greater depth below the water table feels more pressure than does a point at lesser depth. If the water table is horizontal, the pressure acting on an imaginary horizontal reference plane at a specified depth below the water table is the same everywhere. But if the water table is not horizontal, as shown in above, the pressure at points on a horizontal reference plane at depth changes with location. For example, the pressure acting at point p1, which lies below the hill in figure above, is greater than the pressure acting at point p2, which lies below the valley, even though both p1 and p2 are at the same elevation. 

Both the elevation of a volume of groundwater and the pressure within the water provide energy that, if given the chance, will cause the water to flow. Physicists refer to such stored energy as potential energy. The potential energy available to drive the flow of a given volume of groundwater at a location is called the hydraulic head. To measure the hydraulic head at a point in an aquifer, hydrogeologists drill a vertical hole down to the point and then insert a pipe in the hole. The height above a reference elevation (for example, sea level) to which water rises in the pipe represents the hydraulic head water rises higher in the pipe where the head is higher. As a rule, groundwater flows from regions where it has higher hydraulic head to regions where it has lower hydraulic head. This statement generally implies that groundwater regionally flows from locations where the water table is higher to locations where the water table is lower. 

The flow of groundwater.

Hydrogeologists have calculated how hydraulic head changes with location underground, by taking into account both the effect of gravity and the effect of pressure. These calculations reveal that groundwater flows along concave-up curved paths, as illustrated in cross section (figure above a, b). These curved paths eventually take groundwater from regions where the water table is high (under a hill) to regions where the water table is low (below a valley), but because of flow-path shape, 

some groundwater may flow deep down into the crust along the first part of its path and then may flow back up, toward the ground surface, along the final part of its path. The location where water enters the ground (where the flow direction has a downward trajectory) is called the recharge area, and the location where groundwater flows back up to the surface is called the discharge area (see figure above a). 

Flowing water in an ocean current moves at up to 3 km per hour, and water in a steep river channel can reach speeds of up to 30 km per hour. In contrast, groundwater moves at less than a snail’s pace, between 0.01 and 1.4 m per day (about 4 to 500 m per year). Groundwater moves much more slowly than surface water, for two reasons. First, groundwater moves by percolating through a complex, crooked network of tiny conduits, so it must travel a much greater distance than it would if it could follow a straight path. Second, friction between groundwater and conduit walls slows down the water flow. 

Simplistically, the velocity of groundwater flow depends on the slope of the water table and the permeability of the material through which the groundwater is flowing. Thus, groundwater flows faster through high-permeability rocks than it does through low-permeability rocks, and it flows faster in regions where the water table has a steep slope than it does in regions where the water table has a gentle slope. For example, groundwater flows relatively slowly (2 m per year) through a low-permeability aquifer under the Great Plains, but flows relatively quickly (30 m per year) through a high-permeability aquifer under a steep hillslope. In detail, hydrogeologists use Darcy’s Law to determine flow rates at a location.

Darcy’s Law for Groundwater Flow 

 The level to which water rises in a drill hole is the hydraulic head (h). The hydraulic gradient (HG) is the difference in head divided by the length of the flow path.

The rate at which groundwater flows at a given location depends on the permeability of the material containing the groundwater; groundwater flows faster in a more permeable material than it does in a less permeable material. The rate also depends on the hydraulic gradient, the change in hydraulic head per unit of distance between two locations, as measured along the flow path. 

To calculate the hydraulic gradient, we divide the difference in hydraulic head between two points by the distance between the two points as measured along the flow path. This can be written as a formula:

hydraulic gradient = h1 - h2/j

where h1 - h2 is the difference in head (given in meters or feet, because head can be represented as an elevation) between two points along the water table, and j is the distance between the two points as measured along the flow path. A hydraulic gradient exists anywhere that the water table has a slope. Typically, the slope of the water table is so small that the path length is almost the same as the horizontal distance between two points. So, in general, the hydraulic gradient is roughly equivalent to the slope of the water table. 

In 1856, a French engineer named Henry Darcy carried out a series of experiments designed to characterize factors that control the velocity at which groundwater flows between two locations (1 and 2),  each of which has a different hydraulic head (h1 and h2). Darcy represented the velocity of flow by a quantity called the discharge (Q), meaning the volume of water passing through an imaginary vertical plane perpendicular to the groundwater’s flow path in a given time. He found that the discharge depends on the the hydraulic head (h1- h2); the area (A) of the imaginary plane through which the groundwater is passing; and a number called the hydraulic conductivity  (K). The hydraulic conductivity represents the ease with which a fluid can flow through a material. This, in turn, depends on many factors (such as the viscosity and density of the fluid), but mostly it reflects the permeability of the material. The relationship that Darcy discovered, now known as Darcy’s law, can be written in the form of an equation as:

Q = KA(h1 - h2)/j 

The equation states that if the hydraulic gradient increases, discharge increases, and that as conductivity increases, discharge increases. Put in simpler terms, the flow rate of groundwater increases as the permeability increases and as the slope of the water table gets steeper.
Figures credited to Stephen Marshak.

Where Does Groundwater Reside?

Where does groundwater reside?

Groundwater as we know the drinking water which is pulled out of the ground, where does it comes from?

The Underground Reservoir 

Water moves among various reservoirs during the hydrologic cycle. Of the water that falls on land, some evaporates directly back into the atmosphere, some gets trapped in glaciers, and some becomes runoff that enters a network of streams and lakes that drains to the sea. The remainder sinks or percolates downward, by a process called infiltration, into the ground. In effect, the upper part of the crust behaves like a giant sponge that can soak up water.

Of the water that does infiltrate, some descends only into the soil and wets the surfaces of grains and organic material making up the soil. This water, called soil moisture, later evaporates back into the atmosphere or gets sucked up by the roots of plants and transpires back into the atmosphere. But some water sinks deeper into sediment or rock, and along with water trapped in rock at the time the rock formed, makes up groundwater. Groundwater slowly flows underground for anywhere from a few months to tens of thousands of years before returning to the surface to pass once again into other reservoirs of the hydrologic cycle. 

Porosity: Open Space in Rock and Regolith 

Contrary to popular belief, only a small proportion of underground water occurs in caves. Most groundwater resides in  relatively small open spaces between grains of sediment or between grains of seemingly solid rock, or within cracks of various sizes. The term pore refers to any open space within a volume of sediment, or within a body of rock, and the term porosity refers to the total amount of open space within a material, specified as a percentage. For example, if we say that a block of rock has 30% porosity, then 30% of the block consists of pores. Geologists distinguish between two basic kinds of porosity primary and secondary.

Porosity is the open space in rock or sediment, whereas permeability is the degree to which the pores are connected.

Primary porosity develops during sediment deposition and during rock formation (figure above a,b). It includes the pores between clastic grains that exist because the grains don’t fit together tightly during deposition. Secondary porosity refers to new pore space produced in rocks some time after the rock first formed. For example, when rocks fracture, the opposing walls of the fracture do not fit together tightly, so narrow spaces remain in between. Thus, joints and faults may provide secondary porosity for water (figure above c). As groundwater passes through rock, it may dissolve and remove some minerals, creating solution cavities that also provide secondary porosity.

Permeability: The Ease of Flow 

If solid rock completely surrounds a pore, the water in the pore cannot flow to another location. For groundwater to flow, pores must be linked by conduits (openings). The ability of a material to allow fluids to pass through an interconnected network of pores is a characteristic known as permeability. Groundwater flows easily through a material, such as loose gravel, that has high permeability. In gravel, the water is able to pass quickly from pore to pore, so if you pour water into a gravel-filled jar, it will trickle down to the bottom of the jar, where it displaces air and fills the pores (figure above d). In tightly packed sediments or in rock, the water flows more slowly because it follows a tortuous path through tiny conduits. Water flows slowly or not at all through an impermeable material. Put another way, an impermeable material has low permeability or even no permeability. The permeability of a material depends on several factors:

• Number of available conduits: As the number of conduits increases, permeability increases. 
• Size of the conduits: More fluids can travel through wider conduits than through narrower ones. 
• Straightness of the conduits: Water flows more rapidly through straight conduits than it does through crooked ones. 

Note that the factors that control permeability in rock or sediment resemble those that control the ease with which traffic moves through a city. Traffic can flow quickly through cities with many straight, multilane boulevards, whereas it flows slowly through cities with only a few narrow, crooked streets. Porosity and permeability are not the same feature. A material whose pores are isolated from each other can have high porosity but low permeability. 

Aquifers and Aquitards 

Water in the ground-aquifers, aquitards and the water table.

With the concept of permeability in mind, hydrogeologists distinguish between an aquifer, sediment or rock with high permeability and porosity, and an aquitard, sediment or rock that does not transmit water easily and therefore retards the motion of water. An aquifer that is not overlain by an aquitard is an unconfined aquifer. Water can infiltrate down into an unconfined aquifer from the Earth’s surface, and groundwater can rise to reach the Earth’s surface from an unconfined aquifer. An aquifer that is overlain by an aquitard is a confined aquifer its water is isolated from the ground surface (figure above a).

The Water Table 

Infiltrating water can enter permeable sediment and bedrock by percolating along cracks and through conduits connecting pores. Nearer the ground surface, water only partially fills pores, leaving some space that remains filled with air (figure above b). The region of the subsurface in which water only partially fills pores is called the unsaturated zone. Deeper down, water completely fills, or saturates, the pores. This region is the saturated zone. In a strict sense, geologists use the term “groundwater” specifically for subsurface water in the saturated zone, where water completely fills pores. 

The term water table refers to the horizon that separates the unsaturated zone above from the saturated zone below. Typically, surface tension, the electrostatic attraction of water molecules to each other and to mineral surfaces, causes water to seep up from the water table (just as water rises in a thin straw), filling pores in the capillary fringe, a thin layer at the base of the unsaturated zone. Note that the water table forms the top boundary of groundwater in an unconfined aquifer. 

The depth of the water table in the subsurface varies greatly with location. In some places, the water table defines the surface of a permanent stream, lake, or marsh, and thus effectively lies above the ground level (figure above c). Elsewhere, the water table lies hidden below the ground surface. In humid regions, it typically lies within a few meters of the surface, whereas in arid regions, it may lie hundreds of meters below the surface. Rainfall rates affect the water table depth in a given locality  (figure above d) the water table drops during the dry season and rises during the wet season. Streams or ponds that hold water during the wet season may, therefore, dry up during the dry season because their water infiltrates into the ground below.

Topography of the Water Table 

Factors that influence the position of the groundwater.

In hilly regions, if the subsurface has relatively low permeability, the water table is not a planar surface. Rather, its shape mimics, in a subdued way, the shape of the overlying topography (figure above a). This means that the water table lies at a higher elevation beneath hills than it does beneath valleys. But the relief (the vertical distance between the highest and lowest elevations) of the water table is not as great as that of the overlying land, so the surface of the water table tends to be smoother than that of the landscape. 

At first thought, it may seem surprising that the elevation of the water table varies as a consequence of ground-surface topography. After all, when you pour a bucket of water into a pond, the surface of the pond immediately adjusts to remain horizontal. The elevation of the water table varies because groundwater moves so slowly through rock and sediment that it cannot quickly assume a horizontal surface. When rain falls on a hill and water infiltrates down to the water table, the water table rises a little. When it doesn't rain, the water table sinks slowly, but so slowly that when rain falls again, the water table rises before it has had time to sink very far. 

k (such as shale) may lie within a thick aquifer. A mound of groundwater accumulates above such aquitard lenses. The result is a perched water table, a groundwater top surface that lies above the regional water table because the underlying lens of impermeable rock or sediment prevents the groundwater from sinking down to the regional water table (figure above b).
Figures credited to Stephen Marshak.

Coastal Landforms

Where Land Meets Sea: Coastal Landforms

Tourists along the Amalfi coast of Italy thrill to the sound of waves crashing on rocky shores. But in the Virgin Islands sunbathers can find seemingly endless white sand beaches, and along the Mississippi delta, vast swamps border the sea. Large, dome-like mountains rise directly from the sea in Rio de Janeiro, Brazil, but a 100-m-high vertical cliff marks the boundary between the Nullarbor Plain of southern Australia and the Great Southern Ocean. As these examples illustrate, coasts, the belts of land bordering the sea, vary dramatically in terms of topography and associated landforms. 

Beaches and Tidal Flats 

Characteristics of beach, barrier islands and tidal flats.

For millions of vacationers, the ideal holiday includes a trip to a beach, a gently sloping fringe of sediment along the shore. Some beaches consist of pebbles or boulders, whereas others consist of sand grains (figure above a, b). This is no accident, for waves winnow out finer sediment like silt and clay and carry it to quieter water, where it settles. Storm waves, which can smash cobbles against one another with enough force to shatter them, have little effect on sand, for sand grains can’t collide with enough energy to crack. Thus, cobble beaches exist only where nearby cliffs supply large rock fragments. 

The composition of sand varies from beach to beach because different sands come from different sources. Sands derived from the weathering and erosion of silicic-to- intermediate rocks consist mainly of quartz; other minerals in  these rocks chemically weather to form clay, which washes away  in waves. Beaches made from the erosion of limestone, or of  coral reefs and shells, consist of carbonate sand, including masses of sand-sized chips of shells. And beaches derived from  the erosion of basalt boast black sand, made of tiny basalt grains.

A beach profile, a cross section drawn perpendicular to the shore, illustrates the shape of a beach (figure above c). Starting from the sea and moving landward, a beach consists of a foreshore zone, or intertidal zone, across which the tide rises and falls. The beach face, a steeper, concave-up part of the foreshore zone, forms where the swash of the waves actively scours the sand. The backshore zone extends from a small step, cut by high-tide swash to the front of the dunes or cliffs that lie farther inshore. The backshore zone includes one or more berms, horizontal to landward-sloping terraces that receive sediment only during a storm. 

Geologists commonly refer to beaches as “rivers of sand,” to emphasize that beach sand moves along the coast over time it is not a permanent substrate. Wave action at the shore moves an active sand layer on the sea floor on a daily basis. Inactive sand, buried below this layer, moves only during severe storms or not at all. Longshore drift, discussed earlier, can transport sand hundreds of kilometres along a coast in a matter of centuries. Where the coastline indents landward, beach drift stretches beaches out into open water to create a sand spit. Some sand spits grow across the opening of a bay, to form a baymouth bar (figure above d). 

The scouring action of waves sometimes piles sand up in a narrow ridge away from the shore called an offshore bar, which parallels the shoreline. In regions with an abundant sand supply, offshore bars rise above the mean high-water level and become barrier islands (figure above e), and the water between a barrier island and the mainland becomes a quiet-water lagoon, a body of shallow seawater separated from the open ocean. Though developers have covered some barrier islands with expensive resorts, in the time frame of centuries to millennia, barrier islands are temporary features and may wash away in a storm.

Tidal flats, regions of clay and silt exposed or nearly exposed at low tide but totally submerged at high tide, develop in regions protected from strong wave action (figure above f). They are typically found along the margins of lagoons or on shores protected by barrier islands. Here, sediments accumulate to form thick, sticky layers. 

Rocky Coasts 

Erosion landforms of rocky shorelines.

More than one ship has met its end, smashed and splintered in the spray and thunderous surf of a rocky coast, where bedrock cliffs rise directly from the sea. Lacking the protection of a beach, rocky coasts feel the full impact of ocean breakers. The water pressure generated during the impact of a breaker can pick up boulders and smash them together until they shatter, and it can squeeze air into cracks, creating enough force to push rocks apart. Further, because of its turbulence, the water hitting a cliff face carries suspended sand and thus can abrade the cliff. The combined effects of shattering, wedging, and abrading, together called wave erosion, gradually undercut a cliff face and make a wave-cut notch (figure above a). Undercutting continues until the overhang becomes unstable and breaks away at a joint, creating a pile of rubble at the base of the cliff that waves immediately attack and break up. In this process, wave erosion cuts away at a rocky coast, so that the cliff gradually migrates inland. Such cliff retreat may leave behind a wave-cut bench, or platform, that becomes visible at low tide (figure above b). 

Other processes besides wave erosion break up the rocks along coasts. For example, salt spray coats the cliff face above the waves and infiltrates into pores. When the water evaporates, salt crystals grow and push apart the grains, thereby weakening the rock. Biological processes also contribute to erosion, for plants and animals in the intertidal zone bore into the rocks and gradually break them up. 

Many rocky coasts are irregular with headlands protruding into the sea and embayments set back from the sea. Wave energy focuses on headlands and disperses in embayments, a result of wave refraction. The resulting erosion removes debris at headlands, and sediment accumulates in embayments (figure above c). In some cases, a headland erodes in stages (figure above d). Because of refraction, waves curve and attack the sides of a headland, slowly eating through it to create a sea arch connected to the mainland by a narrow bridge. Eventually the arch collapses, leaving isolated sea stacks just offshore (figure above d). Once formed, a sea stack protects the adjacent shore from waves. Therefore, sand may collect in the lee of the stack, slowly building a tombolo, a narrow ridge of sand that links the sea stack to the mainland.

Estuaries 

The Chesapeake Bay estuary formed when the sea flooded river valleys. The region is sinking relative to other coast areas because it overlies a buried meteor crater.

Along some coastlines, a relative rise in sea level causes the sea to flood river valleys that merge with the coast, resulting in estuaries, where seawater and river water mix. You can recognize an estuary on a map by the dendritic pattern of its river-carved coastline (figure above). Oceanic and fluvial waters interact in two ways within an estuary. In quiet estuaries, protected from wave action or river turbulence, the water becomes stratified, with denser oceanic salt water flowing upstream as a wedge beneath less-dense fluvial freshwater.  In turbulent estuaries, oceanic and fluvial water combine to create nutrient-rich brackish water with a salinity between that of oceans and rivers. Estuaries are complex ecosystems inhabited by unique species of shrimp, clams, oysters, worms, and fish that can tolerate large changes in salinity.

Fjords 

Fjord landscapes form where relative sea-level rise drowns glacially carved valleys.

During the last ice age, glaciers carved deep valleys in coastal mountain ranges. When the ice age came to a close, the glaciers melted away, leaving deep, U-shaped valleys. The water stored in the glaciers, along with the water within the vast ice sheets that covered continents during the ice age, flowed back into the sea and caused sea level to rise. The rising sea filled the deep valleys, creating fjords, or flooded glacial valleys. Coastal fjords are fingers of the sea surrounded by mountains; because of their deep-blue water and steep walls of polished rock, they are distinctively beautiful (figure above).

Coastal Wetlands 

Examples of coastal wetlands.

A flat-lying coastal area that floods during high tide and drains during low tide, but does not get pummeled by intense waves, can host salt-resistant plants and evolve into a coastal wetland. Wetland-dominated shorelines are sometimes called “organic coasts.” Researchers distinguish among different types of coastal wetlands based on the plants they host. Examples include swamps (dominated by trees), marshes (dominated by grasses; figure above a), and bogs (dominated by moss and shrubs). So many marine species spawn in wetlands that as a whole, wetlands account for 10% to 30% of marine organic productivity. 

In tropical or semitropical climates (between 30 north and 30 south of the equator), mangrove trees may become the dominant plant in swamps (figure above b). Some mangrove species form a broad network of roots above the water surface, making the plant look like an octopus standing on its tentacles, and some send up small protrusions from roots that rise above the water and allow the plant to breathe. Dense mangrove swamps counter the effects of stormy weather and thus prevent coastal erosion.

Coral Reefs 

The character and evolution of coral reefs.

Along the azure coasts of Hawaii, visitors swim through colorful growths of living coral. Some corals look like brains, others like elk antlers, still others like delicate fans (figure above a). Sea anemones, sponges, and clams grow on and around the coral. Though at first glance coral looks like a plant, it is actually a colony of tiny invertebrates related to jellyfish. An individual coral animal, or polyp, has a tubelike body with a head of tentacles. 

Coral polyps secrete calcite shells, which gradually build into a mound of solid limestone whose top surface lies from just below the low-tide level down to a depth of about 60 m. At any given time, only the surface of the mound lives the mound’s interior consists of shells from previous generations of coral. The realm of shallow water underlain by coral mounds, associated organisms, and debris comprises a coral reef. Reefs absorb wave energy and thus serve as a living buffer zone that protects coasts from erosion. Corals need clear, well-lit, warm (18–30C) water with normal oceanic salinity, so coral reefs grow only along clean coasts at latitudes of less than about 30 (figure above b). 

Marine geologists distinguish three different kinds of coral reef, on the basis of their geometry (figure above c). A fringing reef forms directly along the coast, a barrier reef develops offshore, and an atoll makes a circular ring surrounding a lagoon. As Charles Darwin first recognized back in 1859, coral reefs associated with islands in the Pacific start out as fringing reefs and then later become barrier reefs and finally atolls. This progression reflects the continued growth of the reef as the island around which it formed gradually sinks. Eventually, the reef itself sinks too far below sea level to remain alive and becomes the cap of a flat-topped seamount known as a guyot.

Land drainage

Draining the Land 

Water that drains the land has a series of streams network which is filled from either the ground water or the water from the atmosphere, hydrologic cycle.

Forming Streams and Drainage Networks

Excess surface water (runoff) comes from rain, melting ice or snow, and ground water springs. On flat round, water accumulates in puddles ow swamps, but no slopes, it flows downslope in streams.

Where does the water in a stream come from? Recall that water enters the hydrologic cycle by evaporating from the Earth’s surface and rising into the atmosphere. After a relatively short residence time, atmospheric water condenses and falls back to the Earth’s surface as rain or snow that accumulates in various reservoirs. Some rain or snow remains on the land as surface water (in puddles, swamps, lakes, snowfields, and glaciers), some flows downslope as a thin film called sheetwash, and some sinks into the ground, where it either becomes trapped in soil (as soil moisture) or descends below the water table to become groundwater. (the water table is the level below which groundwater fills all the pores and cracks in subsurface rock or sediment. Above the water table, air partially or entirely fills the pores and cracks.) Streams can receive input of water from all of these reservoirs (figure above). Specifically, gravity pulls surface water (including meltwater) downhill into stream channels, the pressure exerted by the weight of new rainfall squeezes existing soil moisture back out of the ground, and groundwater seeps out of the channel walls into the channel, if the floor of the channel lies below the water table. 

Running water collects in stream channels, because a channel is lower than the surrounding area and gravity always moves material from higher to lower elevation. How does a stream channel form in the first place? The process of channel formation begins when sheetwash starts flowing downslope. Like any flowing fluid, sheetwash erodes its substrate (the material it flows over). The efficiency of such erosion depends on the velocity of the flow faster flows erode more rapidly. In nature, the ground is not perfectly planar, not all substrate has the same resistance to erosion, and the amount of vegetation that covers and protects the ground varies with location. Thus, the velocity of sheetwash also varies with location. Where the flow happens to be a bit faster, or the substrate is a little weaker, erosion scours (digs) a channel. Since the channel is lower than the surrounding ground, sheetwash in adjacent areas starts to head toward it. With time, the extra flow deepens the channel relative to its surroundings, a process called downcutting, and a stream forms. 

 An example of headward erosion. The main stream flows in a deep valley. Side streams are cutting into the bordering plateau.

As its flow increases, a stream channel begins to lengthen at its origin, a process called headward erosion (figure above). Headward erosion occurs for two reasons. First, it happens when the surface flow converging at the entrance to a channel has sufficient erosive power to downcut. Second, it happens at locations where groundwater seeps out of the ground and enters the entrance to the stream channel. Such seepage, called “groundwater sapping,” gradually weakens and undermines the soil or rock just upstream of the channel’s endpoint until the material collapses into the channel; the collapsed debris eventually washes away during a flood. Each increment of collapse makes the channel longer.

As downcutting deepens the main channel, the surrounding land surfaces start to slope toward the channel. Thus, new side channels, or tributaries, begin to form, and these flow into the main channel. Eventually, an array of linked streams evolves, with the smaller tributaries flowing into a trunk stream. The array of interconnecting streams together constitute the drainage network. Like transportation networks of roads, drainage networks of streams reach into all corners of a region, providing conduits for the removal of runoff. 

 Block diagrams illustrating five types of drainage networks.

The configuration of tributaries and trunk streams defines the map pattern of a drainage network. This pattern depends on the shape of the landscape and the composition of the substrate. Geologists recognize several types of networks on the basis of the network’s map pattern (figure above).

• Dendritic: When rivers flow over a fairly uniform substrate with a fairly uniform initial slope, they develop a dendritic network, which looks like the pattern of branches connecting to the trunk of a deciduous tree. 
• Radial: Drainage networks forming on the surface of a cone shaped mountain flow outward from the mountain peak, like spokes on a wheel. Such a pattern defines a radial network. 
• Rectangular: In places where a rectangular grid of fractures (vertical joints) breaks up the ground, channels form along the preexisting fractures, and streams join each other at right angles, creating a rectangular network. 
• Trellis: In places where a drainage network develops across a landscape of parallel valleys and ridges, major tributaries flow down a valley and join a trunk stream that cuts across the ridges. The resulting map pattern resembles a garden trellis, so the arrangement of streams constitutes a trellis network. 
• Parallel: On a uniform slope, several streams with parallel courses develop simultaneously. The group comprises a parallel network.

Drainage Basins and Divides

Drainage divides and basins.

A drainage network collects water from a broad region, variously called a drainage basin, catchment, or watershed, and feeds it into the trunk stream, which carries the water away. The highland, or ridge, that separates one watershed from another is a drainage divide (figure above a, b). A continental divide separates drainage that flows into one ocean from drainage that flows into another. For example, if you straddle the continental divide where it runs along the crest of the Rocky Mountains in the western United States, and pour a cup of water out of each hand, the water in one hand flows to the Atlantic, and the water in the other flows to the Pacific. Three divides bound part of the Mississippi drainage basin, which drains the interior of the United States.

Streams That Last, Streams That Don’t

The contact between permanent and ephemeral streams.

Permanent streams flow all year long, whereas ephemeral streams flow only for part of the year. Some ephemeral streams flow only for tens of minutes to a few hours, following a heavy rain. Most permanent streams exist where the floor (or bed) of the stream channel lies below the water table (figure above a). In these streams, which occur in humid or temperate climates, water comes not only from upstream or from surface runoff, but also from springs through which groundwater seeps. If the bed of a stream lies above the water table, then the stream can be permanent only when the rate at which water  arrives from upstream exceeds the rate at which water infiltrates into the ground below. For example, the downstream portion of the Colorado River in the dry Sonoran Desert of Arizona flows all year, because enough water enters it from the river’s wet headwaters upstream in Colorado; hardly any water enters the stream from the desert itself. 

Streams that do not have a sufficient upstream source, and whose beds lie above the water table, are ephemeral, because the water that fills a channel due to a heavy rain or a spring thaw eventually sinks into the ground and/or evaporates, and the stream dries up (figure above b). Streams whose watersheds lie entirely within an arid region tend to be ephemeral. The dry bed of an ephemeral stream is variously called a dry wash, an arroyo, or a wadi.

Credits: Stephen Marshak (Essentials of Geology)

Numerical age and geologic time

Dating Sedimentary Rocks? 

The mind grows giddy gazing so far back into the abyss of time. John Playfair (1747–1819),  British geologist who popularized the works of Hutton.

We have seen that isotopic dating can be used to date the time when igneous rocks formed and when metamorphic rocks metamorphosed, but not when sedimentary rocks were deposited. So how do we determine the numerical age of a sedimentary rock? We must answer this question if we want to add numerical ages to the geologic column. Geologists obtain dates for sedimentary rocks by studying cross-cutting relationships between sedimentary rocks and datable igneous or metamorphic rocks. For example, if we find a sequence of sedimentary strata deposited unconformably on a datable granite, the strata must be younger than the granite  (figure above). If a datable basalt dike cuts the strata, the strata must be older than the dike. And if a datable volcanic ash buried the strata, then the strata must be older than the ash.

The Geologic Time Scale 

Geologists have searched the world for localities where they can recognize cross-cutting relations between datable igneous  

rocks and sedimentary rocks or for layers of datable volcanic rocks inter-bedded with sedimentary rocks. By isotopically dating the igneous rocks, they have been able to provide numerical ages for the boundaries between all geologic periods. For example, work from around the world shows that the Cretaceous Period began about 145 million years ago and ended 65 million years ago. So the Cretaceous sandstone bed in first figure was deposited during the middle part of the Cretaceous, not at the beginning or end. 

The discovery of new data may cause the numbers defining the boundaries of periods to change, which is why the term numerical age is preferred to absolute age. In fact, around 1995, new dates on rhyolite ash layers above and below the Cambrian-Precambrian boundary showed that this boundary occurred at 542 million years ago, in contrast to previous, less definitive studies that had placed the boundary at 570 million years ago. Figure above shows the currently favoured numerical ages of periods and eras in the geologic column as of 2009. This dated column is commonly called the geologic time scale. 

What Is the Age of the Earth? 

During the 18th and 19th centuries, before the discovery of isotopic dating, scientists came up with a great variety of clever solutions to the question, “How old is the Earth?”—all of which have since been proven wrong. Lord William Kelvin, a 19th century physicist renowned for his discoveries in thermodynamics, made the most influential scientific estimate of the Earth’s age of his time. Kelvin calculated how long it would take for the Earth to cool down from a temperature as hot as the Sun’s, and concluded that this planet is about 20 million years old. Kelvin’s estimate contrasted with those being promoted by followers of Hutton, Lyell, and Darwin, who argued that if the concepts of uniformitarianism and evolution were correct, the Earth must be much older. They argued that physical processes that shape the Earth and form its rocks, as well as the process of natural selection that yields the diversity of species, all take a very long time. Geologists and physicists continued to debate the age issue for many years. The route to a solution didn't appear until 1896, when Henri Becquerel announced the discovery of radioactivity. Geologists immediately realized that the Earth’s interior was producing heat from the decay of radioactive material. This realization uncovered one of the flaws in Kelvin’s argument: Kelvin had assumed that no new heat was produced after the Earth first formed. Because radioactivity constantly generates new heat in the Earth, the planet has cooled down much more slowly than Kelvin had calculated and could be much older. The discovery of radioactivity not only invalidated Kelvin’s estimate of the Earth’s age, it also led to the development of isotopic dating. Since the 1950s, geologists have scoured the planet to identify its oldest rocks. Rocks younger than 3.85 Ga are fairly common. Rock samples from several localities (Wyoming, Canada, Greenland, and China) have yielded dates as old as 4.03 Ga. (Recall that “Ga” means “billion years ago.”) Individual clastic grains of the mineral zircon have yielded dates of up to 4.4 Ga, indicating that rock as old as 4.4 Ga did once exist. Isotopic dating of Moon rocks yields dates of up to 4.50 Ga, and dates on meteorites have yielded ages as old as 4.57 Ga. Geologists consider 4.57-Ga meteorites to be fragments of planetesimals like those from which the Earth first formed. Thus, these dates are close to the age of the Earth’s birth, for models of the Earth’s formation assume that all objects in the Solar System developed at roughly the same time from the same nebula. Why don’t we find rocks with ages between 4.03 and 4.57 Ga in the Earth’s crust? Geologists have come up with several ideas to explain the lack of extremely old rocks. One idea comes from calculations defining how the temperature of our planet has changed over time. These calculations indicate that during the first half-billion years of its existence, the Earth might have been so hot that rocks in the crust remained above the closure temperature for minerals, and isotopic clocks could not start “ticking.” Another idea comes from studies of cratering events on other moons and planets. These studies indicate that the inner planets were bombarded so intensely by meteorites at about 4.0 Ga that almost all crust formed earlier than 4.0 Ga was completely destroyed.

Picturing Geologic Time 

The number 4.57 billion is so staggeringly large that we can’t begin to comprehend it. If you lined up this many pennies in a row, they would make an 87,400-km-long line that would wrap around the Earth’s equator more than twice. Notably, at the scale of our penny chain, human history is only about 100 city blocks long. Another way to grasp the immensity of geologic time is to equate the entire 4.57 billion years to a single calendar year. On this scale, the oldest rocks preserved on Earth date from early February, and the first bacteria appear in the ocean on February 21. The first Shelly invertebrates appear on October 25, and the first amphibians crawl out onto land on November 20. On December 7, the continents coalesce into the super-continent of Pangaea. Birds and the ancestors of mammals  appear about December 15, along with the dinosaurs, and the Age of Dinosaurs ends on December 25. The last week of December represents the last 65 million years of Earth history, including the entire Age of Mammals. The first human-like ancestor appears on December 31 at 3  p.m., and our species, Homo sapiens, shows up an hour before midnight. The last ice age ends a minute before midnight, and all of recorded human history takes place in the last  30 seconds. To put it another way, human history occupies the last 0.000001% of Earth history. The Earth is so old that there has been more than enough time for the rocks and life forms of Earth to have formed and evolved.

How do we determine numerical age of Earth?

Numerical age determination

Geologists since the days of Hutton could determine the relative ages of geologic events, but they had no way to specify numerical ages (called “absolute ages” in older literature). Thus, they could not define a timeline for Earth history or determine the duration of events. This situation changed with the discovery of radioactivity. Simply put, radioactive elements decay at a constant rate that can be measured in the lab and can be specified in years. In the 1950s, geologists developed techniques for using measurements of radioactive elements to calculate the numerical ages of rocks. Geologists originally referred to these techniques as radiometric dating; more recently, this has come to be known as isotopic dating. The overall study of numerical ages is geochronology. Since the 1950s, isotopic dating techniques have steadily improved, and geologists have learned how to make very accurate measurements from very small samples. But the basis of the techniques remains the same, and to explain them, we must first review radioactive decay. 

Radioactive Decay 

All atoms of a given element have the same number of protons in their nucleus we call this number the atomic number. However, not all atoms have the same number of neutrons in their nucleus. Therefore, not all atoms of a given element have the same atomic weight (roughly, the number of protons plus neutrons). Different versions of an element, called isotopes of the element, have the same atomic number but a different atomic weight. For example, all uranium atoms have 92 protons, but the uranium-238 isotope (abbreviated 238U) has an atomic weight of 238 and thus has 146 neutrons, whereas the 235U isotope has an atomic weight of 235 and thus has 143 neutrons. Some isotopes of some elements are stable, meaning that they last essentially forever. Radioactive isotopes are unstable in that eventually, they undergo a change called radioactive decay, which converts them to a different element. Radioactive decay can take place by a variety of reactions that change the atomic number of the nucleus and thus form a different element. In these reactions, the isotope that undergoes decay is the parent isotope, while the decay product is the daughter isotope. For example, rubidium-87 (87Rb) decays to strontium-87 (87Sr), potassium-40 (40K) decays to argon-40 (40Ar), and uranium-238 (238U) decays to lead-206 (206Pb). In some cases, decay takes many steps before yielding a stable daughter. Physicists cannot specify how long an individual radioactive isotope will survive before it decays, but they can measure how long it takes for half of a group of parent isotopes to decay. This time is called the half-life of the isotope. 

Figure above (a-c) can help you visualize the concept of a half-life. Imagine a crystal containing 16 radioactive parent isotopes. (In real crystals, the number of atoms would be much larger.) After one half-life, 8 isotopes have decayed, so the crystal now contains 8 parent and 8 daughter isotopes. After a second half-life, 4 of the remaining parent isotopes have decayed, so the crystal contains 4 parent and 12 daughter isotopes. And after a third half-life, 2 more parent isotopes have  decayed, so the crystal contains 2 parent and 14 daughter isotopes. For a given decay reaction, the half-life is a constant.

Isotopic Dating 

Techniques Since radioactive decay proceeds at a known rate, like the tick-tock of a clock, it provides a basis for telling time. In other words, because an element’s half-life is a constant, we can calculate the age of a mineral by measuring the ratio of parent to daughter isotopes in the mineral. In practice, how can we obtain an isotopic date? First, we must find the right kind of elements to work with. Although there are many different pairs of parent and daughter isotopes among the known radioactive elements, only a few have long enough half-lives, and occur in sufficient abundance in minerals, to be useful for isotopic dating. 

Table above lists particularly useful elements. Each radioactive element has its own half-life. (Note that carbon dating is not used for dating rocks because appropriate carbon isotopes occur only in organisms and radioactive carbon has a very short half-life). Second, we must identify the right kind of minerals to work with. Not all minerals contain radioactive elements, but fortunately some fairly common minerals do. Once we have found the right kind of minerals, we can set to work using the following steps. 

• Collecting the rocks: We need to find un-weathered rocks for dating, for the chemical reactions that happen during weathering may lead to the loss of some isotopes. 
• Separating the minerals: The rocks are crushed, and the appropriate minerals are separated from the debris. 
• Extracting parent and daughter isotopes: To separate out the parent and daughter isotopes from minerals, we can use several techniques, including dissolving the minerals in acid or evaporating portions of them with a laser. 
• Analyzing the parent-daughter ratio: Once we have a sample of appropriate atoms, we pass them through a mass spectrometer, an instrument that uses a strong magnet to separate isotopes from one another according to their respective weights (figure below). The instrument can count the number of atoms of specific isotopes separately. 

At the end of the laboratory process, we can define the ratio of parent to daughter isotopes in a mineral, and from this ratio calculate the age of the mineral. Needless to say, the description of the procedure here has been simplified in reality, obtaining an isotopic date is time-consuming and expensive and requires complex calculations.

What Does an Isotopic Date Mean? 

At high temperatures, atoms in a crystal lattice vibrate so rapidly that chemical bonds can break and reattach relatively easily. As a consequence, isotopes can escape from or move into crystals, so parent-daughter ratios are meaningless. Because isotopic dating is based on the parent-daughter ratio, the “isotopic clock” starts only when crystals become cool enough for isotopes to be locked into the lattice. The temperature below which isotopes are no longer free to move is called the closure temperature of a mineral. When we specify an isotopic date for a mineral, we are defining the time at which the mineral cooled below its closure temperature. With the concept of closure temperature in mind, we can interpret the meaning of isotopic dates. In the case of igneous rocks, isotopic dating tells you when a magma or lava cooled to form a solid, cool igneous rock. In the case of metamorphic rocks, an isotopic date tells you when a rock cooled from a metamorphic temperature above the closure temperature to a temperature below. Not all minerals have the same closure temperature, so different minerals in a rock that cools very slowly will yield different dates. Can we isotopically date a clastic sedimentary rock directly? No. If we date minerals in a sedimentary rock, we determine only when these minerals first crystallized as part of an igneous or metamorphic rock, not the time when the minerals were deposited as sediment nor the time when the sediment lithified to form a sedimentary rock. For example, if we date the feldspar grains contained within a granite pebble in a conglomerate, we’re dating the time the granite cooled below feldspar’s closure temperature, not the time the pebble was deposited by a stream.

Other Methods of Determining Numerical Age

The rate of tree growth depends on the season. During the spring, trees grow rapidly and produce lighter, less-dense wood, but during the winter trees grow slowly or not at all, and produce darker, denser wood. Thus, wood contains recognizable annual growth rings. Such tree rings provide a basis for determining age. If you've ever wondered how old a tree that’s just been cut down might be, just look at the stump and count the rings. Notably, by correlating clusters of distinctive rings in the older parts of living trees with comparable clusters of rings in dead logs, scientists can extend the tree-ring record back for many thousands of years, allowing geologists to track climate changes back into prehistory. Seasonal changes also affect rates of such phenomena as shell growth, snow accumulation, clastic sediment deposition, chemical sediment precipitation, and production of organic material. Geologists have learned to use growth rings in shells, as well as rhythmic layering in sediments and in glacial ice (figure above a–c), to date events numerically back through recent Earth history.