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WEATHERING
Physical weathering
Physical weathering of rocks is the breakdown of rocks into smaller size particles by pure mechanical processes without changing the chemical composition and mineralogy, except for the removal of some soluble components due to erosion. Many sedimentary rocks are composed of particles that have been weathered, eroded, transported, and terminally deposited in basins. Sandstone is formed from bonded
sand-sized particles under water. Its porosity makes it vulnerable to the processes of physical weathering. Physical weathering reduces the particle size and compactness, and increases the surface area and bulk volume.
Exfoliation is the process of peeling the outer layers of the rock from the main body due to differential expansion and contraction between the outer and interior mass of the rock. It occurs in areas where there is extreme variation between day and night temperatures of the order of 25°C-30°C, for example in deserts. Exfoliation is also sometimes called spheroidal weathering, when spherical boulders are formed due to smoothening of sharp edges due to exfoliation. Exfoliation is also known as insolation weathering or thermal insulation. Differential expansion and contraction may also occur at constant temperature due to the variation in the colors of mineral grains in rock.
Frost Wedging
The physical disintegration of rocks by the wedging action of ice is called frost wedging. In polar regions, water is frozen in the form of ice. During the day time, ice melts due to higher temperature, percolates, and fills the cracks and fissures existing in the rocks.
Mineral Crystallization
The process of mineral crystallization is similar to that of frost wedging. When salts – occurring in the form of solution in rock fissures and cracks – undergo crystallization, the volume of salt minerals increases, causing pressure on the walls of the fissures and cracks. The increase in temperature also causes expansion of the salts, resulting in additional pressure on the walls of the fissures and cracks, leading to disintegration of the rock.
Slaking
Slaking is alternate wetting and drying of rocks and can be a very important factor in weathering. Slaking occurs by the mechanism of “ordered water,” which is the accumulation of successive layers of water molecules in between the mineral grains of a rock.
Action of Vegetation
Growth of vegetation in rocky terrain, causes the roots of trees and plants to enlarge and extend through weak planes of the rocks. This leads to widening the existing cracks and creating new cracks in the rock mass.
Chemical Weathering
Chemical weathering is the decomposition of rocks by a change in the chemical and mineralogical composition, through a combination of several chemical processes. It is a slow but more intense process than physical weathering. Most of the chemical weathering processes occur in the presence of water. Chemical weathering takes place mainly at the surface of rocks and minerals, leading to disappearance of certain minerals and formation of new products and secondary minerals.
During the process of chemical weathering, one or more of the following components are formed: Minerals in solution (cations and anions).
b. Oxides of iron and alumina (sesquioxides Al2O3, Fe2O3).
c. Various forms of silica (silicon oxide compounds).
d. Stable wastes such as very fine silt (mostly fine quartz) and sand (coarser quartz).
Carbonation
Carbonation is the process in which the carbonic acid and other acids are responsible for chemical weathering. Carbonic acid (H2CO3) is formed when carbon dioxide in the atmosphere dissolves in rain water, as shown by the following chemical reaction – CO2 + H2O → H2CO3
Carbonation is the process in which carbonic acid reacts with the calcium carbonate in rocks and forms calcium bicarbonate, which is soluble in water – CaCO3 + H2CO3 → Ca (HCO3)2
Solution
Solution is one of the processes of chemical decomposition of rocks in which the water dissolves and removes soluble cementing materials such as calcium carbonate. When rocks are continually exposed to water or subject to action of water over long duration, the water soluble substances are removed from the rock. The rock no longer remains solid and forms holes or rills, and ultimately breaks into pieces or decomposes. Most of the rock minerals are affected by this process including limestone, marl, calcareous shale, dolomite, quartz, etc.
Hydrolysis
Hydrolysis is the most important process in chemical weathering. It occurs due to the dissociation of water molecules (H2O) into hydrogen (H+) and hydroxyl (OH)– ions, which chemically combine with minerals and bring about changes, such as ion-exchange, decomposition of crystalline structure, and formation of new compounds.
Hydration
Hydration is the adsorption of water on rock surface. It generally precedes all other processes in chemical weathering. When the rock is in contact with water for a long duration, the disintegration of rock takes place due to hydration. Soil-forming minerals in rocks undergo hydration (wetting with water), when exposed to humid conditions.
Oxidation
Oxidation is the process in which the oxygen ions combine with the minerals in rocks, causing the removal of one or more electrons from a compound. This weakens the mineral structure and makes it less rigid and unstable, causing decomposition of minerals. Oxidation of rocks is similar in process to the corrosion of steel. Iron oxides formed by oxidation give the red color to the red soil. The most common oxides are those of iron and aluminum, and their respective red and yellow staining of soils is quite common in tropical regions, which have high temperatures and precipitation.
Reduction
It is the process of removal of oxygen and is the reverse of oxidation. It is equally important in changing soil color to gray, blue, or green as ferric iron is converted to ferrous iron compounds. Reduction takes place under the conditions of excess water or waterlogged condition with little or no oxygen.
Complexation
Metals released from primary minerals such as Fe, Mn, and Al build complexes with organic components, such as fulvic acid and humic acid, which are very stable. Weathering of primary minerals produces secondary minerals. Elements released from primary minerals are prone to leaching if they do not form complexes.
EARTH’S INTERIOR
We know through seismography that temperatures in the inner parts of the Earth actually exceed the surface temperature of the Sun! That is pretty hot. So why is the center of the Earth Hot. The answer comes from a lot different sources. The first is heat left over from the formation of the Earth. The next source is gravitational pressure put on core by tidal forces and the rotation of the Earth. The last known source of heat is the radioactive decay of elements in the inner part of the Earth.
The Earth is pretty old at 4 billion years old and there are still things we don’t completely understand about its formation. We do know that gravity played a role pulling in more matter and compressing it to form the Earth. When you have matter colliding at high velocities like it did in the early stages of the Solar System’s development all that kinetic energy has to go somewhere. In the case of Earth that energy was turned into heat. This heat is the initial source for the temperatures in the Earth’s interior.
The next source of heat is gravitational pressure. The Earth is under immense pressure due to the tidal forces exerted by the Sun, the Moon, and the other planets in the Solar System. When you include the fact that it is also rotating the Earth’s core is under immense pressure. This pressure basically keeps the core hot in the same way as a pressure cooker. It also helps to minimize the heat it loses.
The last and most important source of heat is nuclear fission of heavy elements in the Earth’s interior. In short, the Earth has a nuclear engine inside it. It is thank to the continuous nuclear fission of elements in the Earth’s interior that replaces the heat the Earth loses keeping it nice and hot. This fission process occurs in the form of radioactive decay. It also creates the convection currents in the mantle that drive plate tectonics.
MAGMA
Magma is a molten and semi-molten rock mixture found under the surface of the Earth. This mixture is usually made up of four parts: a hot liquid base, called the melt; minerals crystallized by the melt; solid rocks incorporated into the melt from the surrounding confines; and dissolved gases. When magma is ejected by a volcano or other vent, the material is called lava. Magma that has cooled into a solid is called igneous rock.
How Magma Forms
Earth is divided into three general layers. The core is the superheated center, the mantle is the thick, middle layer, and the crust is the top layer on which we live. Magma originates in the lower part of the Earth’s crust and in the upper portion of the mantle. Most of the mantle and crust are solid, so the presence of magma is crucial to understanding the geology and morphology of the mantle. Differences in temperature, pressure, and structural formations in the mantle and crust cause magma to form in
different ways.
Decompression Melting
Decompression melting involves the upward movement of
Earth's mostly-solid mantle. This hot material rises to an area of lower pressure through the process of convection. Areas of lower pressure always have a lower melting point than areas of high pressure. This reduction in overlying pressure, or decompression, enables the mantle rock to melt and form magma. Decompression melting often occurs at divergent boundaries, where tectonic plates separate.
Transfer of Heat
Magma can also be created when hot, liquid rock intrudes into Earth’s cold crust. As the liquid rock solidifies, it loses its heat to the surrounding crust. Much like hot fudge being poured over cold ice cream, this transfer of heat is able to melt the surrounding rock into magma. Transfer of heat often happens at convergent boundaries, where tectonic plates are crashing together. As the denser tectonic plate subducts, or sinks below, or the less dense tectonic plate, hot rock from below can intrude into the cooler plate above. This process transfers heat and creates magma.
Flux Melting
Flux melting occurs when water or carbon dioxide are added to rock. These compounds cause the rock to melt at lower temperatures. This creates magma in places where it originally maintained a solid structure. Much like heat transfer, flux melting also occurs around subduction zones. In this case, water overlying the subducting seafloor would lower the melting temperature of the mantle, generating magma that rises to the surface.
Magma Escape Routes
Magma leaves the confines of the upper mantle and crust in two major ways: as an intrusion or as an extrusion. An intrusion can form features such as dikes and xenoliths. An extrusion could include lava and volcanic rock. Magma can intrude into a low-density area of another geologic formation, such as a sedimentary rock structure. When it cools to solid rock, this intrusion is often called a pluton. A pluton is an intrusion of magma that wells up from below the surface. Plutons can include dikes and xenoliths. A magmatic dike is simply a large slab of magmatic material that has intruded into another rock body. A xenolith is a piece of rock trapped in another type of rock. Many xenoliths are crystals torn from inside the Earth and embedded in magma while the magma was cooling. The most familiar way for magma to escape, or extrude, to Earth’s surface is through lava. Lava eruptions can be “fire fountains” of liquid rock or thick, slow-moving rivers of molten material. Lava cools to form volcanic rock as well as volcanic glass. Magma can also extrude into Earth’s atmosphere as part of a violent volcanic explosion. This magma solidifies in the air to form volcanic rock called tephra. In the atmosphere, tephra is more often called volcanic ash. As it falls to Earth, tephra includes rocks such as pumice.
Magma Chamber
In areas where temperature, pressure, and structural formation allow, magma can collect in magma chambers. Most magma chambers sit far beneath the surface of the Earth. The pool of magma in a magma chamber is layered. The least-dense magma rises to the top. The densest magma sinks near the bottom of the chamber. Over millions of years, many magma chambers simply cool to form a pluton or large igneous intrusion.
If a magma chamber encounters an enormous amount of pressure, however, it may fracture the rock around it. The cracks, called fissures or vents, are tell-tale signs of a volcano. Many volcanoes sit over magma chambers. As a volcano’s magma chamber experiences greater pressure, often due to more magma seeping into the chamber, the volcano may undergo an eruption. An eruption reduces the pressure inside the magma chamber. As long as more magma pools into a volcano’s magma chamber, there is the possibility of an eruption and the volcano will remain active. Large eruptions can nearly empty the magma chamber. The layers of magma may be documented by the type of eruption material the volcano emits. Gases, ash, and light-colored rock are emitted first, from the least-dense, top layer of the magma chamber. Dark, dense volcanic rock from the lower part of the magma chamber may be released later. In violent eruptions, the volume of magma shrinks so much that the entire magma chamber collapses and forms a caldera.
Types of Magma
Mafic Magma
Mafic magma has relatively low silica content, roughly 50%, and higher contents in iron and magnesium. This type of magma has a low gas content and low viscosity, or resistance to flow. Mafic magma also has high mean temperatures, between 1000o and 2000oCelsius (1832o and 3632o Fahrenheit), which contributes to its lower viscosity. Low viscosity means that mafic magma is the most fluid of magma types. It erupts non-explosively and moves very quickly when it reaches Earth’s surface as lava.
Intermediate Magma
Intermediate magma has higher silica content (roughly 60%) than mafic magma. This results in a higher gas content and viscosity. Its mean temperature ranges from 800o to 1000o Celsius (1472o to 1832oFahrenheit). As a result of its higher viscosity and gas content, intermediate magma builds up pressure below the Earth’s surface before it can be released as lava. This more gaseous and sticky lava tends to explode violently and cools as andesite rock.
Felsic Magma
Felsic magma has the highest silica content of all magma types, between 65-70%. As a result, felsic magma also has the highest gas content and viscosity, and lowest mean temperatures, between 650o and 800o Celsius (1202o and 1472o Fahrenheit). Thick, viscous felsic magma can trap gas bubbles in a volcano’s magma chamber. These trapped bubbles can cause explosive and destructive eruptions. These eruptions eject lava violently into the air, which cools into dacite and rhyolite rock.
Definition of Metamorphism
The word "Metamorphism" comes from the Greek: Meta = change, Morph = form, so metamorphism means to change form. In geology this refers to the changes in mineral assemblage and texture that result from subjecting a rock to pressures and temperatures different from those under which the rock originally formed. The original rock that has undergone metamorphism is called the protolith. Protolith can be any type of rock and sometimes the changes in texture and mineralogy are so dramatic that it is difficult to distinguish what the protolith was.
Factors that Control Metamorphism
Metamorphism occurs because rocks undergo changes in temperature and pressure and may be subjected to differential stress and hydrothermal fluids. Metamorphism occurs because some minerals are stable only under certain conditions of pressure and temperature. When pressure and temperature change, chemical reactions occur to cause the minerals in the rock to change to an assemblage that is stable at the new pressure and temperature conditions. But, the process is complicated by such things as how the pressure is applied, the time over which the rock is subjected to the higher pressure and temperature, and whether or not there is a fluid phase present during metamorphism.
• Temperature
Temperature increases with depth in the Earth along the Geothermal Gradient. Thus, higher temperature can occur by burial of rock. Temperature can also increase due to igneous intrusion.
• Pressure increases with depth of burial, thus, both pressure and temperature will vary with depth in the Earth. Pressure is defined as a force acting equally from all directions. It is a type of stress, called hydrostatic stress, or uniform stress. If the stress is not equal from all directions, then the stress is called a differential stress.
• Fluid Phase - Any existing open space between mineral grains in a rock can potentially contain a fluid. This fluid is mostly H2O, but contains dissolved ions. The fluid phase is important because chemical reactions that involve changing a solid mineral into a new solid mineral can be greatly speeded up by having dissolved ions transported by the fluid. If chemical alteration of the rock takes place as a result of these fluids, the process is called metasomatism.
•Time - Because metamorphism involves changing the rock while it is solid, metamorphic change is a slow process. During metamorphism, several processes are at work. Recrystallization causes changes in minerals size and shape. Chemical reactions occur between the minerals to form new sets of minerals that are more stable at the pressure and temperature of the environment, and new minerals form as a result of polymorphic phase transformations (recall that polymorphs are compounds with the same chemical formula, but different crystal structures.
There are two kinds of differential stress. Normal stress causes objects to be compressed in the direction of maximum principal stress and extended in the direction of minimal stress. If differential stress is present during metamorphism, it can have a profound effect on the texture of the rock. Shear stress causes objects to be smeared out in the direction of applied stress. Differential stress if acting on a rock can have a profound effect on the appearance or texture of the rock.
STRESS
Causes and Types of Stress
Stress is the force applied to an object. In geology, stress is
the force per unit area that is placed on a rock. Four types of
stresses act on materials.
• A deeply buried rock is pushed down by the weight of all
the material above it. Since the rock cannot move, it cannot
deform. This is called confining stress.
• Compression squeezes rocks together, causing rocks to fold or fracture (break) (Figure below). Compression is the most common stress at convergent plate boundaries.
Stress caused these rocks to fracture.
• Rocks that are pulled apart are under tension. Rocks under tension lengthen or break apart. Tension is the major type of stress at divergent plate boundaries.
• When forces are parallel but moving in opposite directions, the stress is called shear (Figure below). Shear stress is the most common stress at transform plate boundaries. When stress causes a material to change shape, it has undergone strain or deformation.
Rocks have three possible responses to increasing stress:
• elastic deformation: the rock returns to its original shape when the stress is removed. • plastic deformation: the rock does not return to its original shape when the stress is removed. • fracture: the rock breaks.
• At the Earth's surface, rocks usually break quite quickly, but deeper in the crust, where temperatures and pressures are higher, rocks are more likely to deform plastically.
• Sudden stress, such as a hit with a hammer, is more likely to make a rock break. Stress applied over time often leads to plastic deformation.
SEAFLOOR SPREADING
Seafloor spreading is a process that occurs at mid-ocean ridges, where new oceanic crust is formed through volcanic activity and then gradually moves away from the ridge. Seafloor spreading helps explain continental drift in the theory of plate tectonics. When oceanic plates diverge, tensional stress causes fractures to occur in the lithosphere. Basaltic magma rises up the fractures and cools on the ocean floor to form new sea floor. Older rocks will be found farther away from the spreading zone while younger rocks will be found nearer to the spreading zone.
Incipient spreading
In the general case, sea floor spreading starts as a rift in a continental land mass, similar to the Red Sea-East Africa Rift System today. The process starts with heating at the base of the continental crust which causes it to become more plastic and less dense. Because less dense objects rise in relation to denser objects, the area being heated becomes a broad dome. As the crust bows upward, fractures occur that gradually grow into rifts. The typical rift system consists of three rift arms at approximately 120-degree angles. These areas are named triple junctions and can be found in several places across the world today. The separated margins of the continents evolve to form passive margins. Hess' theory was that new seafloor is formed when magma is forced upward toward the surface at a mid-ocean ridge. If spreading continues past the incipient stage described above, two of the rift arms will open while the third arm stops opening and becomes a 'failed rift'.
Continued spreading and subduction
As new seafloor forms and spreads apart from the mid-ocean ridge it slowly cools over time. Older seafloor is therefore colder than new seafloor, and older oceanic
basins deeper than new oceanic basins due to isostasy.
If the diameter of the earth remains relatively constant
despite the production of new crust, a mechanism must
exist by which the crust is also destroyed. The destruction of
oceanic crust occurs at subduction zones where
oceanic crust is forced under either continental crust or
oceanic crust.
OCEAN BASIN
Most of the earth's surface is covered by oceans, but for
a long time the oceans have been an essentially white
spot on the map of the world. Early expeditions like that
of the Beagle (Charles Darwin) brought some preliminary
knowledge, compilations of data by ship captains
brought some initial knowledge about ocean currents
and migration of fish swarms (mention Melville, Captain Ahab), but by and far we did not know much about the topography of the ocean floor, much less about its geological features. Starting at around 1930, however, a vast amount of knowledge has been gathered about the oceans, about their water chemistry, the cycling of elements, biological aspects, bathymetry, bottom sediments and their stratigraphy.
Though much less spectacular and not as well publicized, the progress in knowledge about the oceans is far more important for the future of mankind than to send a few men to the moon. Ocean research has implications for food resources, the supply of raw materials for a growing population, and possibilities of ocean population by man (giant raft cities in shallow seas, platforms moving with food-rich ocean currents, etc.). Even populating the deep sea is probably cheaper and more feasible than to have people live in colonies on the moon.
Work on the bathymetry of the ocean basins (mainly with echo-sounding devices) has revealed many morphologic features that were previously unknown, such as oceanic ridges, abyssal plains (and hills), seamounts, trenches, and continental margins, all of these features are now easily explained by plate tectonics.
Continental Shelf = flooded edges of the continents; Continental Margin = the edge/border region of the continent; Deep Sea Trenches = deepest parts of ocean basins (due to subduction of oceanic crust); Abyssal Plains= older parts of oceanic crust, smoothed due to sediment deposition; Seamounts = submarine volcanic cones; the can also form linear arrangements, so called Seamount Chains.
Continental margins are in a geological sense not part of the oceanic crust. They consist of continental crust and material that was eroded from the continents and is now piled up along the margins of the continents. The margins are subdivided into CONTINENTAL SLOPE and SHELF with the latter simply being a submerged part of shield or platform.
PLATE BOUNDARIES
Types of plate boundaries
Three types of plate boundaries exist, with a fourth, mixed type, characterized by the way the plates move relative to each other. They are associated with different types of surface phenomena. The different types of plate boundaries are:
Transform boundary Divergent boundary Convergent boundary
Transform boundaries (Conservative) occur where two lithospheric plates slide, or perhaps more accurately, grind past each other along transform faults, where plates are neither created nor destroyed. Transform faults occur across a spreading center. Strong earthquakes can occur along a fault.
Divergent boundaries (Constructive) occur where two plates slide apart from each other. At zones of ocean-to ocean rifting, divergent boundaries form by seafloor spreading, allowing for the formation of new ocean basin. As the ocean plate splits, the ridge forms at the spreading center, the ocean basin expands, and finally, the plate area increases causing many small volcanoes and/or shallow earthquakes.
Convergent boundaries (Destructive) (or active margins) occur where two plates slide toward each other to form either a subduction zone (one plate moving underneath the other) or a continental collision. At zones of ocean to-continent subduction, the dense oceanic lithosphere plunges beneath the less dense continent.
Deep marine trenches are typically associated with subduction zones, and the basins that develop along the active boundary are often called "foreland basins". Closure of ocean basins can occur at continent-to-continent boundaries (e.g., Himalayas and Alps): collision between masses of granitic continental lithosphere; neither mass is subducted; plate edges are compressed, folded, uplifted. Plate boundary zones occur where the effects of the interactions are unclear, and the boundaries, usually occurring along a broad belt, are not well defined and may show various types of movements in different episodes.
STRATIFIED ROCKS
The Structure Of Rock Masses - Stratified Rocks
The stratified rocks form more than nine-tenths of the earth's surface, and if the entire series of them were present at any one place, they would have a maximum thickness of about thirty miles, but no such place is known. The regions of greatest sedimentary accumulation are the shallower parts of the oceans, while those regions which have remained as dry land, through long ages, may not only have had no important additions to their surfaces, but have lost immense thicknesses of rock through denudation. The great oceanic abysses are also areas of excessively slow sedimentation, and thus the thickness of the stratified rocks varies much from point to point, a variation which has been increased by the irregularities of upheaval and depression and of different rates of denudation.
Stratification, or division into layers, is the most persistent and conspicuous characteristic of the sedimentary rocks. In studying the sedimentary deposits of the present day, we learned that by the sorting power of water and wind, heterogeneous material is arranged into more or less homogeneous beds, separated from one another by distinct planes of division, and the same thing is true of the sedimentary rocks of all ages. This division into more or less parallel layers is called stratification, and the extent to which the division is carried varies according to circumstances.
A single member, or bed, of a stratified rock, whether thick or thin, is called a layer, though for purposes of distinction, excessively thin layers are called lamince. Each layer or lamina represents an uninterrupted deposition of material, while the divisions between them, or bedding planes, are due to longer or shorter pauses in the process, or to a change, if only in a film, of the material deposited.
RELATIVE AND ABSOLUTE DATING
Absolute dating
Geologists often need to know the age of material that they find. They use absolute dating methods, sometimes called numerical dating, to give rocks an actual date, or date range, in number of years. This is different to relative dating, which only puts geological events in time order.
Radiometric dating
Most absolute dates for rocks are obtained with radiometric methods. These use radioactive minerals in rocks as geological clocks. The atoms of some chemical elements have different forms, called isotopes. These break down over time in a process scientists call radioactive decay. Each original isotope, called the parent, gradually decays to form a new isotope, called the daughter. Each isotope is identified with what is called a ‘mass number’.
Absolute dating rock layers
Isotopes are important to geologists because each radioactive element decays at a constant rate, which is unique to that element. These rates of decay are known, so if you can measure the proportion of parent and daughter isotopes in rocks now, you can calculate when the rocks were formed. Because of their unique decay rates, different elements are used for dating different age ranges.
Radiocarbon dating measures radioactive isotopes in once-living organic material instead of rock, using the decay of carbon-14 to nitrogen-14. Because of the fairly fast decay rate of carbon-14, it can only be used on material up to about 60,000 years old. Geologists use radiocarbon to date such materials as wood and pollen trapped in sediment, which indicates the date of the sediment itself.
What is an isotope?
Measuring isotopes is particularly useful for dating igneous and some metamorphic rock, but not sedimentary rock. Sedimentary rock is made of particles derived from other rocks, so measuring isotopes would date the original rock material, not the sediments they have ended up in. However, there are radiometric dating methods that can be used on sedimentary rock, including luminescence dating. All radiometric dating methods measure isotopes in some way. Most directly measure the amount of isotopes in rocks, using a mass spectrometer. Others measure the subatomic particles that are emitted as an isotope decays. Some measure the decay of isotopes more indirectly.
Relative dating
Relative dating is used to arrange geological events, and the rocks they leave behind, in a sequence. The method of reading the order is called stratigraphy (layers of rock are called strata). Relative dating does not provide actual numerical dates for the rocks.
GEOLOGIC TIME
Fossils and relative dating
Fossils are important for working out the relative
ages of sedimentary rocks. Throughout the
history of life, different organisms have
appeared, flourished and become extinct.
Many of these organisms have left their remains
as fossils in sedimentary rocks. Geologists have
studied the order in which fossils appeared and
disappeared through time and rocks. This study
is called biostratigraphy. Fossils can help to
match rocks of the same age, even when you
find those rocks a long way apart. This matching
process is called correlation, which has been an
important process in constructing geological
timescales. Some fossils, called index fossils, are
particularly useful in correlating rocks. For a fossil
to be a good index fossil, it needs to have lived
during one specific time period, be easy to
identify and have been abundant and found in
many places.
Geologic Time
Geologists and paleontologists usually represent geologic time vertically. This arrangement is derived from the vertical succession of rock strata in Earth itself. Because of the Principle of Superposition, geologists know that new layers can be laid down only on top of pre-existing older strata, and therefore older rocks lie below younger ones. The vertical time scale mimics this stratigraphic arrangement by placing older time periods below younger ones, with the present day at the very top. By itself, this vertical arrangement reveals only the relative ages of rocks, not how old the layers are.
Many different rocks from different strata around the world have been dated this way. Once numerical ages have been obtained, they can be combined with the vertical time succession to create a true geologic time scale. This works because some of the strata that have been dated are the same ones that were used to create the relative vertical time scale. The result is a global, comprehensive time scale that lists all the ages of Earth’s history, with the oldest ages placed at the bottom and the youngest at the top.
Major Divisions of Geologic Time
The major divisions, with brief explanations of each, are shown in the following scale of relative geologic time, which is arranged in chronological order with the oldest division at the bottom, the youngest at the top.
While every fossil tells us something about the age of the rock it's found in, index fossils are the ones that tell us the most. Index fossils (also called key fossils or type fossils) are those that are used to define periods of geologic time.
A good index fossil is one with four characteristics: it is distinctive, widespread, abundant and limited in geologic time. Because most fossil-bearing rocks formed in the ocean, the major index fossils are marine organisms. That being said, certain land organisms are useful in young rocks and in specific regions.
Any type of organism can be distinctive, but not so many are widespread. Many important index fossils are of organisms that start out life as floating eggs and infant stages, which allowed them to populate the world using ocean currents. The most successful of these became abundant, yet at the same time, they became the most vulnerable to environmental change and extinction. Thus, their time on Earth may have been confined to short period of time.
Consider trilobites, a very good index fossil for Paleozoic rocks that lived in all parts of the ocean. Trilbotes were a class of animal, just like mammals or reptiles, meaning that the individual species within the class had noticeable differences. Trilobites were constantly evolving new species during their existence, which lasted 270 million years from Middle Cambrian time to the end of the Permian Period, or almost the entire length of the Paleozoic. Because they were mobile animals, they tended to inhabit large, even global areas. They were also hard-shelled invertebrates, so they fossilized easily. These fossils are large enough to study without a microscope.
Other index fossils of this type include ammonites, crinoids, rugose corals, brachiopods, bryozoans and mollusks. The USGS offers a more detailed list of invertebrate fossils (with scientific names only) here.
Other major index fossils are small or microscopic, part of the floating plankton in the world ocean. These are handy because of their small size. They can be found even in small bits of rock, such as wellbore cuttings. Because their tiny bodies rained down all over the ocean, they can be found in all kinds of rocks. Therefore, the petroleum industry has made great use of index microfossils, and geologic time is broken down in quite fine detail by various schemes based on graptolites, fusulinids, diatoms and radiolarians.
The rocks of the ocean floor are geologically young, as they are constantly subducted and recycled into the Earth's mantle. Thus, marine index fossils older than ~200 million years are normally found in sedimentary strata on land, in areas that were once covered by seas.
For terrestrial rocks, which form on land, regional or continental index fossils may include small rodents that evolve quickly as well as larger animals that have wide geographic ranges. These form the basis of provincial time divisions.
Index fossils are used in the formal architecture of geologic time for defining the ages, epochs, periods and eras of the geologic time scale. Some of the boundaries of these subdivisions are defined by mass extinction events, like the Permian-Triassic extinction. The evidence for these events is found in the fossil record wherever there is a disappearance of major groups of species within a geologically short amount of time.
Related fossil types include the characteristic fossil—a fossil that belongs to a time period but doesn't define it— and the guide fossil, one that helps narrow down a time range rather than nail it down. Geologists have organized the history of the Earth into a timescale on which large chunks of time are called periods and smaller ones are called epochs.
Index Fossils
Keyed to the relative time scale are examples of index fossils, the forms of life which existed during limited periods of geologic time and thus are used as guides to the age of the rocks in which they are preserved.
The first people who needed to understand the geological relationships of different rock units were miners. Mining had been of commercial interest since at least the days of the Romans, but it wasn't until the 1500s and 1600s that these efforts produced an interest in local rock relationships. By noting the relationships of different rock units, Nicolaus Steno in 1669 described two basic geologic principles. The first stated that sedimentary rocks are laid down in a horizontal manner, and the second stated that younger rock units were deposited on top of older rock units. To envision this latter principle think of the layers of paint on a wall. The oldest layer was put on first and is at the bottom, while the newest layer is at the top. An additional concept was introduced by James Hutton in 1795, and later emphasized by Charles Lyell in the early 1800s. This was the idea that natural geologic processes were uniform in frequency and magnitude throughout time, an idea known as the "principle of uniformitarianism".
Steno's principles allowed workers in the 1600s and early 1700s to begin to recognize rock successions. However, because rocks were locally described by the color, texture, or even smell, comparisons between rock sequences of different areas were often not possible. Fossils provided the opportunity for workers to correlate between geographically distinct areas. This contribution was possible because fossils are found over wide regions of the earth's crust. For the next major contribution to the geologic time scale we turn to William Smith, a surveyor, canal builder, and amateur geologist from England. In 1815 Smith produced a geologic map of England in which he successfully demonstrated the validity of the principle of faunal succession. This principle simply stated that fossils are found in rocks in a very definite order. This principle led others that followed to use fossils to define increments within a relative time scale.
The oldest known glacial period is the Huronian. Based on evidence of glacial deposits from the area around Lake Huron in Ontario and elsewhere, it is evident that the Huronian Glaciation lasted from approximately 2,400 to 2,100 Ma. Because rocks of that age are rare, we don’t know much about the intensity or global extent of this glaciation. A succession of incredibly harsh ice ages waxed and
waned during the Cryogenian Period wherein the glaciation was so severe that it was possible to have reached even the equator. Late in the Proterozoic, for reasons that are not fully understood, the climate cooled dramatically and Earth was seized by what appears to be its most intense glaciation. There ere two main glacial periods within the Cryogenian, each lasting for about 20 million years: the Sturtian at around 700 Ma and the Marinoan at 650 Ma. There is also evidence of some shorter glaciations both before and after these. The end of the Cryogenian glaciations coincides with the evolution of relatively large and complex life forms on Earth. This started during the Ediacaran Period, and then continued with the so-called explosion of life forms.