Inside the Planet

Inside the Planet

kacimo samy   3:51 ص

Introduction

Seismologists, geologists who study seismic waves noticed in the early 20th century that P waves bended, or refracted, in their journey through Earth. Observations at stations far removed from the earthquake focus recorded waves that had traveled through the planet’s interior, as illustrated in part (1) of the figure.. Travel times of these waves indicated a refracted path, as shown in the figure, and wave speed is the distance divided by time (as determined by the amount of time elapsed since the start of the earthquake). Refraction was not too surprising because the increased pressure in Earth’s interior results in firmer structures and more resistance to oscillation, so the wave speed is greater and seismic waves refract. What surprised early seismologists was that beyond a certain point about 7,200 miles (11,600 km) from the focus, at an angular distance of 105 degrees S waves disappeared! 

In 1906 the British seismologist Richard D. Oldham (1858–1936) proposed that the disappearance of the shear waves was due to the “shadow” of a liquid core. Since S waves are shear, they cannot propagate through liquid, so the existence of a liquid center inside the planet would explain why seismometers fail to record shear waves on the other side of the planet from the focus, as shown in part (2) of the fi gure below. P waves, being compression waves, refract at the boundary between rock and liquid, creating a smaller “shadow.” Th e rocky interior beneath the crust is called the mantle, and in 1914 the German seismologist Beno Gutenberg (1889–1960) used the seismic wave results to calculate that the mantlecore boundary is located at a depth of about 1,800 miles (2,900 km) below the surface. However, in 1936 the Danish seismologist Inge Lehmann (1888– 1993) analyzed seismic wave data and discovered an additional refractory step of P waves. Her analysis suggested the existence of another boundary, which she placed at a depth of about 3,200 miles (5,150 km). This boundary is between an outer core and an inner core. 

The use of seismic waves to image Earth’s interior is similar to the use of ultrasound waves to image the body’s interior or sound waves in sonar to image the seafloor. Unlike ultrasound and sonar techniques, though, seismologists usually do not generate seismic waves these are natural occurrences beyond the control of researchers. Yet the waves reveal a lot of information about otherwise inaccessible places. Seismic waves are also plentiful; about 1 million or so earthquakes occur each year in the world, and although most of these are fortunately minor they are detectable with sensitive instruments. 

By studying the nature and speed of seismic waves, geologists have learned much about the Earth’s interior. Earth consists of the following several layers:

• crust, composed of rocks having relatively low density, extending from the continental surface to an average depth of about 22 miles (35 km) and from the ocean floor an average of about four miles (6.4 km) down to a boundary known as the Mohorovicic discontinuity (Moho for short), named after the Croatian scientist Andrija Mohorovičić (1857–1936); 
• mantle, extending from the crust to about 1,800 miles (2,900 km) below the surface, and divided into an upper and a lower section; 
• outer core, which is liquid and extends from the mantle border to a depth of about 3,200 miles (5,150 km); 
• inner core, which is solid, with a radius of about 750 miles (1,220 km).

The mantle gets its name from Wiechert, who thought of it as a coat that covered the core (mantle derives from the German word, mantel, for “shell” or “coat”). About 67 percent of Earth’s mass is contained in this large region. The mantle is mostly solid, although as discussed below there is some degree of fluidity in spots; it consists of minerals such as olivine and another silicate called perovskite (MgSiO3). Silicon and aluminium are less abundant in the mantle compared to the crust, but magnesium is much more plentiful. 

Wiechert assumed from the studies of Earth’s density that the core must be dense. A greater density for the core also makes sense because the large portion of the heavier elements would have sunk to the interior as the hot, molten planet formed long ago. Iron and nickel possess relatively high densities and are commonly found in certain meteorites, indicating their abundance throughout the solar system. These metals are likely constituents of the core. The absence of shear wave propagation indicates the outer core is liquid, but studies of other seismic waves indicates a density slightly less than that expected if the outer core contained only melted iron and nickel. Instead, the outer core is about 90 percent iron and nickel, and most of this is iron about 85 percent of the outer core is made of this element. The remaining 10 percent consists of lighter elements such as sulphur and oxygen.

The inner core forms a boundary with the outer core, reflecting some of the waves and transmitting the rest. Shear waves cannot pass through the outer core, but as compression waves cross the boundary between the inner and outer core, some of these disturbances create shear waves. The shear waves travel through the inner core and get converted back into compression waves as they proceed from the inner to the outer core. Seismologists can detect the paths of these waves, and the propagation of shear waves in the inner core implies it cannot be liquid. Density studies suggest the inner core is mostly solid iron, mixed with a small percentage of nickel. 

Researchers continue to study seismic waves and similar data to learn more of the details on the structure and composition inside Earth. In 2005 John W. Hernlund and Paul J. Tackley of the University of California, Los Angeles, and Christine Thomas of the University of Liverpool in the Britain found data suggesting the presence of a thin layer around the mantle-core boundary. This layer, previously unknown and not yet widely studied, might help scientists to understand and identify further properties of the mantle. The researchers published their report “A Doubling of the Post-Perovskite Phase Boundary and Structure of the Earth’s Lowermost Mantle” in a 2005 issue of Nature.

Although researchers can study the finer structure of Earth’s hidden interior with sensitive seismometers, a large amount of information could also be gained by burrowing inside and taking a look. There are limitations on how far down people can drill, even with the hardest bits (the tip of the drill), but researchers are sharpening their drill bits in the effort to reach greater depths.

Petroleum system

Petroleum system

kacimo samy   4:40 ص

Petroleum System

In order to understand the petroleum generation and extraction petroleum system should be known
Petroleum system starts with the deposition to storage from where the production is obtained.
Petroleum system journey starts with the deposition of organic matter.

Deposition of organic matter

The deposition of organic matter starts when organism starts to die and deposits deep down the ocean floor and the above deposition of clay (finer grains). The clay particles are about 1/256 mm size and is called shale. Organic matter deposited on the ocean floor cannot be oxidized due to the depth factor so they can produce hydrocarbon. Hydrocarbon generation needs the cooking of organic matter at high temperature and pressure and it is obtained when it goes into overburden of deposition by clay particles and greater depths. 

Source rock

In the petroleum system the source rock are the shale (clay that goes under high pressure and temperature which cooks the organic matter). Sometimes limestone can also be the source rock with 1% of organic matter contains. So theses rocks undergo cooking where the temperature and pressure determines what type of fuel will be generated. Despite of temperature and pressure another factor in producing hydrocarbon is the time span required to generate fuel. The time is a critical factor as if the organic matter is cooked for less time it will not generate hydrocarbon and when it is greater than the oil produce will be converted into gas.

Reservoir rock

Reservoir as indicated by the name reserves of hydrocarbon. the hydrocarbon cannot be obtained from the source rock because of the higher pores but are lesser to none interconnection. For the extraction the pores should be interconnected so that it can travel when are extracted. But if there is no reservoir and obtaining fuel from source rock it must be fractured for permeability generation. Reservoir rock are mostly sandstone which have higher porosity and permeability but in some cases limestone also serves as reservoir rock. Limestone all by self is not a good reservoir due to fine particles present which give less permeability but as limestone is calcium carbonate so it can be dissolved in water which are the Karst topography. Only then can limestone have permeability required for hydrocarbon to be obtained.

Migration

Primary migration

Migration itself is cleared so the primary migration occurs when hydrocarbon moves from source rock to the reservoir rocks. Primary migration occurs when the source rock is fractured due to tectonic forces (plate movements) or by the overburden squeezing the source rock. As HC (hydrocarbon) have low density they moves upward. 

Secondary migration

Secondary migration is the HC movement within the reservoir rocks. The HC will moves upward in the reservoir rocks.

Seal rock

Trap or seal rocks are those that are present above reservoir rocks as HC movement will always be upward. Seal rock are those that have low to none permeability so that HC cannot escape but are trapped within the reservoir rocks. Shale can be seal rock also as they have porosity but do not have permeability factor so HC will be trapped. Types of traps include stratigraphic and structural.
Stratigraphic traps example is shale as a seal rock and structural traps are fold or faults. 

Time period

The last thing in the petroleum system is time as have said it above already, time is required for HC generation which is always critical. No more time and no less time while cooking of the organic matter or it will not produce the fuel.

The Pleistocene Ice Age

The Pleistocene Ice Age

kacimo samy   10:47 م

The Pleistocene Ice Age 

The Pleistocene Glaciers 

Today, most of the land surface in New York City lies hidden beneath concrete and steel, but in Central Park it’s still possible to see land in a seminatural state. If you stroll through the park, you’ll find that the top surfaces of outcrops are smooth and polished, and in places have been grooved and scratched. Here and there, glacial erratics rest on the bedrock. You are seeing evidence that an ice sheet once scraped along this now-urban ground. Geologists estimate that the ice sheet that overrode the New York City area may have been 250 m thick, enough to bury a 75-story building. 

The fact that glaciated landscapes still decorate the surface of the Earth means that the last ice age occurred fairly recently during Earth’s history. Otherwise, the landscape features would have been either eroded away or buried. The ice age responsible for the glaciated landscapes of North America, Europe, and Asia happened mostly during the Pleistocene Epoch, which began about 2.6 Ma, so as we’ve noted earlier, it is commonly known as the Pleistocene Ice Age. 

Pleistocene ice sheets of the northern hemisphere.

Based on studying patterns of glacial striations and of the sources of erratics, geologists have developed an approximate idea of where the great Pleistocene ice sheets originated, and the directions in which the ice sheets flowed. In North America, major ice sheets appear to have initiated in at least three locations (figure above). The Labrador ice sheet formed over north-eastern Canada, the Keewatin ice sheet originated in north-western Canada, and the Baffin ice sheet formed over 

Baffin Bay. These sheets, together with one or more smaller ones, merged to form the giant Laurentide ice sheet that covered all of Canada east of the Rocky Mountains, and spread southward over the northern portion of the United States. The Cordilleran ice sheet, which originated in the mountains of western Canada, spread westward to the Pacific coast and eastward until it merged with the Laurentide ice sheet. Other ice sheets formed in Greenland, Scandinavia, northern Russia, and Siberia.

In addition to continental ice sheets, sea ice in the  northern hemisphere expanded to cover all of the Arctic Ocean and parts of the North Atlantic during the Pleistocene. Sea ice surrounded Iceland and approached Scotland and also fringed most of western Canada and southeastern Alaska. 

Life and Climate in the Pleistocene World 

Climate belts during the Pleistocene.

During the Pleistocene Ice Age, all climatic belts shifted southward (figure above a, b). Geologists can document this shift by examining fossil pollen, which can survive for thousands of years if preserved in the sediment of bogs. 

Fossils also tell us that numerous species of now-extinct large mammals inhabited the Pleistocene world (figure above c). Giant mammoths and mastodons, relatives of the elephant, along with woolly rhinos, musk oxen, reindeer, giant ground sloths, bison, lions, saber-toothed cats, giant cave bears, and hyenas wandered forests and tundra in North America. Early human-like species were already foraging in the woods by the beginning of the Pleistocene Epoch, and by the end modern Homo sapiens lived on every continent except Antarctica, and had discovered fire and invented tools. 

Timing of the Pleistocene Ice Age 

Louis Agassiz assumed that only one ice age had affected the planet. But close examination of the stratigraphy of glacial deposits on land revealed that paleosols (ancient soil preserved in the stratigraphic record), as well as beds containing fossils of warmerweather animals and plants, lay between distinct layers of glacial sediment. This observation suggested that between episodes of glacial deposition, glaciers receded and temperate climates prevailed. In the second half of the 20th century, when modern methods for dating geological materials became available, the difference in ages between the different layers of glacial sediment could be confirmed. Clearly, glaciers advanced and then retreated more than once during the Pleistocene. Times during which the glaciers grew and covered substantial areas of the continents are called glacial periods, or glaciations, and times between glacial periods are called interglacial periods, or interglacials. 

Pleistocene glacial deposits in the north-central United States. Curving moraines reflect the shape of glacial lobes.

Using the on-land sedimentary record, geologists recognized five Pleistocene glaciations in Europe and, traditionally, four in the mid-western United States (Wisconsinan, Illinoian, Kansan, and Nebraskan, named after the southernmost states in which their till was deposited; figure above). Since the mid-1980s, geologists no longer recognize Nebraskan and Kansan; they are lumped together as “pre-Illinoian.” 

The chronology of glaciations was turned on its head in the 1960s, when geologists began to study submarine sediment. They found that some layers contained glacially transported grains, while others did not. Similarly, they found that at a given location, some layers contained fossils of cold-water plankton and other fossils of warm-water plankton. Researchers found that in sediment of the last 2.6 million years, there is evidence for 20 to 30 glaciations during the Pleistocene Epoch. The traditionally recognized glaciations of Europe and the United States might represent only the largest of these. 

The timing of glaciations. Ice ages have occurred at several times in the geologic past.

Geologists refined their conclusions about the frequency of Pleistocene glaciations by examining the isotopic composition of fossil shells. Shells of many plankton species consist of calcite (CaCO3). The oxygen in the shells includes two isotopes, a heavier one (18O) and a lighter one (16O). The ratio of these isotopes tells us about the water temperature in which the plankton grew; this is because as water gets colder, plankton incorporate a higher proportion of 18O into their shells. The isotope record confirms that 20 to 30 of these events occurred during the last 2.6 million years (figure above a). 

Older Ice Ages during Earth History 

So far, we’ve focused on the Pleistocene Ice Age because of its importance in developing Earth’s present landscape. Was this the only ice age during Earth history, or do ice ages happen frequently? To answer such questions, geologists study the stratigraphic record and search for ancient glacial deposits that have hardened into rock. These deposits, called tillites, consist of larger clasts distributed throughout a matrix of sandstone and mudstone. In many cases, tillites are deposited on glacially polished surfaces.

By using the stratigraphic principles, geologists have determined that tillites were deposited during the Late Paleozoic; these are the deposits Alfred Wegener studied when he argued in favour of continental drift (figure above b). Tillites were also deposited between 850 and 630 Ma (at the end of the Proterozoic Eon), about 2.4 to  2.1 Ga (near the beginning of the Proterozoic), and perhaps about 2.9 Ga (in the Archean Eon). Strata deposited at other times in Earth history do not contain tillites. Thus, it appears that glacial advances and retreats have not occurred steadily throughout Earth history, but rather are restricted to specific time intervals, or ice ages, of which there were four or five: Pleistocene, Permian, late Proterozoic, early Proterozoic, and perhaps Archean. 

Of particular note, some tillites of the late Proterozoic event were deposited at equatorial latitudes, suggesting that, for at least a short time, the continents worldwide were largely glaciated, and the sea may have been covered worldwide by ice. Geologists refer to the ice-encrusted planet as snowball Earth.
Credits: Stephen Marshak (Essentials of Geology)

Unconformities: Gaps in the Record

Unconformities: Gaps in the Record

kacimo samy   5:58 ص

Unconformities: Gaps in the Record 

The unconformity at Siccar Point, Scotland.

To find good exposures of rock, James Hutton sometimes boated along the coast of Scotland, where waves of the stormy North Sea have stripped away soil and shrubbery. He was particularly puzzled by an outcrop at Siccar Point, where two distinct sequences of sedimentary rock lie in contact (figure above a, b). In the lower portion of the outcrop, beds of gray sandstone and shale dip nearly vertically, whereas in the upper portion, beds of red sandstone and conglomerate display a dip of less than 20 degree. Further, the gently dipping layers seem to lie across the truncated ends of the vertical layers, like a handkerchief lying across a row of books. We can imagine that as Hutton was examining this odd geometric relationship, the tide came in and deposited a new layer of sand on top of the rocky shore. With the principle of uniformitarianism in mind, Hutton suddenly realized the significance of what he saw. The gray sandstone-shale sequence had been deposited, turned into rock, tilted, and truncated by erosion before the red sandstone–conglomerate beds had been deposited.

Hutton deduced that the surface between the gray and red rock sequences represented a time interval during which new strata had not been deposited at Siccar Point and the older strata had been eroded away. We now call such a surface, representing a period of non-deposition and possibly erosion, an unconformity. The gap in the geologic record that an unconformity represents is called a hiatus. Geologists recognize three kinds of unconformities: 

The three kinds of unconformities and their formation.

• Angular unconformity: Rocks below an angular unconformity were tilted or folded before the unconformity developed (figure above a). Thus, an angular unconformity cuts across the underlying layers, and the orientation of layers below an unconformity differs from that of the layers above. (The outcrop at Siccar Point exposes an angular unconformity.) 
• Nonconformity: A nonconformity is a type of unconformity at which sedimentary rocks overlie generally much older intrusive igneous rocks and/or metamorphic rocks (figure above b). The igneous or metamorphic rocks underwent cooling, uplift, and erosion prior to becoming the substrate, or basement, on which new sediments accumulated. 
• Disconformity: Imagine that a sequence of sedimentary beds has been deposited beneath a shallow sea. Then sea level drops, exposing the beds for some time. During this time, no new sediment accumulates, and some of the pre-existing sediment gets eroded away. Later, sea level rises, and a new sequence of sediment accumulates over the old. The boundary between the two sequences is a disconformity (figure above c, d). Even though the beds above and below the disconformity are parallel, the contact between them represents an interruption in deposition.

The succession of strata at a particular location provides a record of Earth history there. But because of unconformities, the record preserved in the rock layers is incomplete. It’s as if geologic history is being chronicled by a tape recorder that turns on only intermittently when it’s on (times of deposition), the rock record accumulates, but when it’s off (times of non-deposition and possibly erosion), an unconformity develops. Because of unconformities, no single location on Earth contains a complete record of Earth history.
Credits: Stephen Marshak (Essentials of Geology)

Rock Deformation

Rock Deformation

kacimo samy   5:55 ص

Rock Deformation 

What Are Deformation and Strain? 

Deformation changes the character and configuration of rocks.

To get a visual sense of what geologists mean by the term deformation, let’s contrast rock that has not been affected by an orogeny with rock that has been affected. Our “undeformed” example comes from a road cut in the Great Plains of North America, and our “deformed” example comes from a cliff in the Alps of Europe (figure above a, b).

The road cut, which lies at an elevation of only 100 m above sea level, exposes nearly horizontal beds of sandstone, shale, and limestone these beds have the same orientation that they had when first deposited. Sand grains in the sandstone of this outcrop have a nearly spherical shape (the same shape they had when deposited), and clay flakes in the shale lie roughly parallel to the bedding (because of compaction). When we say that rock of this outcrop is undeformed, we mean that it contains no geologic structures other than a few joints. 

Rocks of the Alpine cliff, exposed at an elevation of 3 km, look very different. Here, we find layers of quartzite, slate, and marble (the metamorphic equivalents of sandstone, shale, and limestone). The layers are cut by joints, but they have also been contorted into folds. On closer examination, we also note that the grains in the quartzite aren't spherical, like those of a sandstone, but have been flattened so that they are somewhat elliptical. And in the slate, clay flakes are aligned to form a type of foliation called slaty cleavage. Finally, if we try tracing the quartzite and slate layers along the outcrop face, we might find that they abruptly terminate at a sloping surface marked by broken-up rock. This surface is a fault. In our example, thick layers of marble lie below the fault the quartzite and slate must have moved along the fault from where they first formed to get to their present location juxtaposed against the marble. 

The components of deformation include displacement, rotation, and distortion.

e Alpine cliff emphasize that during deformation, rocks can undergo one or more of the following (figure above a–c): (1) a change in location (displacement); (2) a change in orientation (rotation); and (3) a change in shape (distortion). Geologists commonly refer to distortion as strain. We distinguish among different kinds of strain according to the nature of the shape change developed during deformation. If a layer of rock becomes longer, it has undergone stretching, but if the layer becomes shorter, it has undergone shortening (figure belowa–c). And, if a change in shape involves the movement of one part of a rock body past another so that angles between features in the rock change, the result is called shear strain (figure below d, e).

Different kinds of strain in rock. Strain is a measure of the distortion, or change in shape, that takes place in rock during deformation.

Brittle versus Ductile Deformation 

Brittle versus ductile deformation.

Imagine that a plate tumbles off a table and lands on a hard floor the plate breaks and smashes into pieces. Similarly, if you strike a glass window with a hammer, the window cracks and may even shatter. Such cracking and breaking serve as familiar examples of brittle deformation (figure above a, b). Now, imagine that you squeeze a ball of soft dough between a book and a tabletop the dough flattens into a pancake. Similarly, if you bend a stick of chewing gum, it changes from a plane into a curve. During such ductile deformation, objects change shape without visibly breaking (figure above c, d). 

What actually happens within mineral grains during these two different kinds of deformation? Recall that the atoms that make up mineral grains are connected by chemical bonds. During brittle deformation, many bonds break and stay broken, leading to the formation of a permanent crack across which material no longer connects. During ductile deformation, some bonds break but new ones quickly form. In this way, the atoms within grains rearrange, and the grains change shape without permanent cracks forming. 

Why do rocks inside the Earth sometimes deform brittlely and sometimes ductilely? The behavior of a rock depends on:

• Temperature: Warmer rocks tend to deform ductilely, whereas colder rocks tend to deform brittlely. Heat makes materials softer. 
• Pressure: Under great pressures deep in the Earth, rock behaves more ductilely than it does under low pressures near the surface. Pressure effectively prevents rock from s eparating into fragments. 
• Deformation rate: A sudden change in shape causes brittle  deformation, whereas a slow change in shape causes ductile deformation. For example, if you hit a marble bench with a hammer, it shatters, but if you leave the bench alone for a century, it gradually sags without breaking. 
• Composition: Some rock types are softer than others; for example, halite (rock salt) deforms ductilely under conditions in which granite deforms brittlely. 

Considering that pressure and temperature both increase with depth in the Earth, geologists find that, in typical continental crust, rocks generally behave brittlely above a depth of about 10 to 15 km, and ductilely below this depth. The depth at which this change in behaviour takes place is called the brittle-ductile transition. Earthquakes in continental crust happen only above this depth because these earthquakes involve brittle breaking.

Stress and the Causes of Deformation 

Up to this point, we've focused on picturing the consequences of deformation. Describing the causes of deformation is a bit more challenging. In captions for displays about mountain building, museums and national parks typically dispense with the issue by using the phrase “The mountains were caused by forces deep within the Earth.” But what does this mean? Isaac Newton stated that a force can cause an object to speed up, slow down, or change direction. In the context of geology, plate interactions and continent-continent  collisions apply forces to rock and thus cause rock to change location, orientation, or shape. In other words, the application of forces in the Earth indeed causes deformation.

There are several kinds of stress.

Geologists, however, use the word stress instead of force when talking about the cause of deformation. We define the stress acting on a plane as the force applied per unit area of the plane. The need to distinguish between stress and force arises because the actual consequences of applying a force depend not just on the amount of force but also on the area over which the force acts. A pair of simple experiments shows why (figure above a). Experiment 1: Stand on a single, empty aluminium can. All of your weight a force focuses entirely on the can, and the can crushes. Experiment 2: Place a board atop 100 cans and stand on the board. In this case, your weight is distributed across 100 cans, and the cans don’t crush. In both experiments, the force caused by the weight of your body was the same, but in Experiment 1 the force was applied over a small area so a large stress developed, whereas in Experiment 2 the same force was applied over a large area so only a small stress developed. How does this concept apply to geology? During mountain building, the force of one plate interacting with another is distributed across the area of contact between the two plates, so the deformation resulting at any specific location actually depends on the stress developed at that location, not on the total force produced by the plate interaction.

Different kinds of stress occur in rock bodies (figure above b–e). Compression takes place when a rock is squeezed, tension occurs when a rock is pulled apart, and shear stress develops when one part of a rock body moves sideways past another. Pressure refers to a special stress condition that happens when the same push acts on all sides of an object. 

Note that stress and strain have very different meanings to geologists, even though we tend to use the words interchangeably in everyday English: stress refers to the amount of force applied per unit area of a rock, whereas strain refers to a change in shape of a rock. Thus, stress causes strain. Specifically, compression causes shortening, tension leads to stretching, and shear stress produces shear strain. Pressure can cause an object to become smaller, but will not cause it to change shape. With our knowledge of stress and strain, we can now look at the nature and origin of various classes of geologic structures.

Credits: Stephen Marshak (Essentials of Geology)

Sedimentary Basins

Sedimentary Basins

kacimo samy   7:38 ص

Sedimentary Basins

The sedimentary veneer on the Earth’s surface varies greatly in thickness. If you stand in central Siberia or south-central Canada, you will find yourself on igneous and metamorphic basement rocks that are over a billion years old sedimentary rocks are nowhere in sight. Yet if you stand along the southern coast of Texas, you would have to drill through over 15 km of sedimentary beds before reaching igneous and metamorphic basement. Thick accumulations of sediment form only in special regions where the surface of the Earth’s lithosphere sinks, providing space in which sediment collects. Geologists use the term subsidence to refer to the process by which the surface of the lithosphere sinks, and the term sedimentary basin for the sediment-filled depression. In what geologic settings do sedimentary basins form? An understanding of plate tectonics theory provides the answers.

Categories of Basins in the Context  of Plate Tectonics Theory 

The geologic setting of sedimentary basins.

Geologists distinguish among different kinds of sedimentary basins in the context of plate tectonics theory. Let’s consider a few examples (figure above). 

1. Rift basins: These form in continental rifts, regions where the lithosphere is stretching horizontally, and therefore thins  vertically. As the rift grows, slip on faults drops blocks of crust down, producing low areas bordered by narrow mountain ridges. These troughs fill with sediment. 
2. Passive-margin basins: These form along the edges of continents that are not plate boundaries. They are underlain by stretched lithosphere, the remnants of a rift whose evolution successfully led to the formation of a mid-ocean ridge and subsequent growth of a new ocean basin. Passive-margin  basins form because subsidence of stretched lithosphere  continues long after rifting ceases. They fill with sediment carried to the sea by rivers and with carbonate rocks formed in coastal reefs. 
3. Intracontinental basins: These develop in the interiors of continents, initially because of subsidence over a rift. They  may continue to subside in pulses even hundreds of millions of years after they formed, for reasons that are not well  understood. 
4. Foreland basins: These form on the continent side of a mountain belt because the forces produced during convergence or collision push large slices of rock up faults and onto the surface of the continent. The weight of these slices pushes down on the surface of the lithosphere, producing a wedge-shaped depression adjacent to the mountain range that fills with sediment eroded from the range. Fluvial and deltaic strata accumulate in foreland basins.

Transgression and Regression 

Sea-level changes, relative to the land surface, control the succession of sediments that we see in a sedimentary basin. At times during Earth history, sea level has risen by as much as a couple of hundred meters, creating shallow seas that submerge the interiors of continents. At other times, sea level has fallen by a couple of hundred meters, exposing the continental shelves to air. Global sea-level changes may be due to a number of factors, including climate change, which controls the amount of ice stored in polar ice caps and causes changes in the volume of ocean basins. Sea level at a location may also be due to the local uplift or sinking of the land surface.

The concept of transgression and regression, during deposition of sedimentary sequence.

When relative sea level rises, the shoreline migrates inland. We call this process transgression. When relative sea level falls, the coast migrates seaward. We call this process regression (figure above). The process of transgression and regression leads to the formation of broad blankets of sediment.

Diagenesis 

Earlier we discussed the process of lithification, by which sediment hardens into rock. Lithification is an aspect of a broader phenomenon called diagenesis. Geologists use the term diagenesis for all the physical, chemical, and biological processes that transform sediment into sedimentary rock and that alter characteristics of sedimentary rock after the rock has formed. 

In sedimentary basins, sedimentary rocks may become very deeply buried. As a result, the rocks endure higher pressures and temperatures and come in contact with warm groundwater. Diagenesis, under such conditions, can cause chemical reactions in the rock that produce new minerals and can also cause cement to dissolve or precipitate.

As temperature and pressure increase still deeper in the subsurface, the changes that take place in rocks become more profound. At sufficiently high temperature and pressure, metamorphism begins, in that a new assemblage of minerals forms, and/or mineral grains become aligned parallel to each other. The transition between diagenesis and metamorphism in sedimentary rocks is gradational and occurs between temperatures of 150C and 300C. In the next chapter, we enter the realm of true metamorphism.

Credits: Stephen Marshak (Essentials of Geology)

Movement and Solidification of Molten Rock

Movement and Solidification of Molten Rock

kacimo samy   2:05 ص

Movement and Solidification of Molten Rock 

If magma stayed put once it formed, new igneous rocks would not develop in or on the crust. But it doesn't stay put; magma tends to move upward, away from where it formed. In some cases, it reaches the Earth’s surface and erupts at a volcano. This movement is a key component of the Earth System, because it transfers material from deeper parts of the Earth upward and provides the raw material from which new rocks and the atmosphere and ocean form. Eventually, magma freezes and transforms into a new solid rock.

Why Does Magma Rise? 

Magma rises for two reasons. First, buoyancy drives magma upward just as it drives a wooden block up through water, because magma is less dense than the surrounding rock. Second, magma rises because the weight of overlying rock creates pressure at depth that literally squeezes magma upward. The same process happens when you step into a puddle barefoot and mud squeezes up between your toes.

What Controls the Speed of Flow? 

Viscosity affects lava behaviour .

Viscosity, or resistance to flow, affects the speed with which magmas or lavas move. Magmas with low viscosity flow more easily than those with high viscosity, just as water flows more easily than molasses. Viscosity depends on temperature, volatile content, and silica content. Hotter magma is less viscous than cooler magma, just as hot tar is less viscous than cool tar, because thermal energy breaks bonds and allows atoms to move more easily. Similarly, magmas or lavas containing more volatiles are less viscous than dry (volatile-free) magmas, because the volatiles also tend to break apart silicate molecules and may also form gas bubbles. Mafic magmas are less viscous than felsic magmas, because silicon-oxygen tetrahedra tend to link together in magma to create long molecular chains that can’t move past each other easily, and there are more of these chains in a felsic magma than in a mafic magma. Thus, hotter mafic lavas have relatively low viscosity and flow in thin sheets over wide regions, but cooler felsic lavas are highly viscous and may clump into a dome-like mound at the volcanic vent (figure above a, b). 

Transforming Melt into Rock 

If a melt stayed at its point of origin, and nothing in its surroundings changed, it would stay molten. But melts don’t last forever. Rather, they eventually solidify or “freeze.” This process happens, in some cases, because volatiles escape from the melt, so that the freezing temperature rises if the melt’s temperature stays the same but its freezing temperature rises, it will solidify. Most often, however, freezing takes place simply when melt cools below its freezing temperature. Temperature decreases upward, toward the Earth’s surface, so magma enters a cooler environment automatically as it rises. If it is trapped underground as an intrusion, it slowly loses heat to surrounding wall rock, drops below its freezing temperature, and solidifies. If melt extrudes as lava at the ground surface, it cools in contact with air or water. 

The time it takes for a magma to cool depends on how fast it is able to transfer heat into its surroundings. To see why, think about the process of cooling coffee. If you pour hot coffee into a thermos bottle and seal it, the coffee stays hot for hours; because it’s insulated, the coffee in the thermos loses heat to the air outside only very slowly. Like the thermos bottle, surrounding wall rock acts as an insulator in that it transports heat away from a magma only very slowly, so magma underground (in an intrusive environment) cools slowly. In contrast, if you spill coffee on a table, it cools quickly because it loses heat to the cold air. Similarly, lava that erupts at the ground surface cools quickly because the air or water surrounding it can conduct heat away quickly. 

Three factors control the cooling time of magma that freezes below the surface in the intrusive realm. 

Factors that affect the freezing of molten rocks.

• The depth of intrusion: Magma intruded deep in the crust, where it is surrounded by warm wall rock, cools more slowly than does magma intruded into cold wall rock near the ground surface. 
• The shape and size of a magma body: Heat escapes from magma at an intrusion’s surface, so the greater the surface area for a given volume of intrusion, the faster it cools. Thus, a body of magma roughly with the shape of a pancake cools faster than one with the shape of a melon. And since the ratio of surface area to volume increases as size decreases, a body of magma the size of a car cools faster than one the size of a ship (figure above a, b). 
• The presence of circulating groundwater: Water passing through magma absorbs and carries away heat, much like the coolant that flows around an automobile engine.

Changes in Magma during Cooling:  Fractional Crystallization 

Most people are familiar with the process of forming ice out of liquid water cool the water to a temperature of 0nC and crystals of ice start to form. Keep the temperature cold enough for long enough and all the water becomes solid, composed entirely of one type of mineral water ice. The process of freezing magma or lava is much more complex, because molten rock contains many different compounds, not just water, so during freezing of molten rock, many different minerals form. Further, not all of these minerals form at the same time (Bowen’s Reaction Series). To get a sense of this complexity, let’s look at an example.

When a mafic magma starts to freeze, mafic (iron- and magnesium-rich) minerals such as olivine and pyroxene start to crystallize first. These solid crystals are denser than the remaining magma, so they start to sink (figure above c). Some react chemically with the remaining magma as they sink, but some reach the floor of the magma chamber and become isolated from the magma. This process of sequential crystal formation and settling is called fractional crystallization it progressively extracts iron and magnesium from the magma, so the remaining magma becomes more felsic. If a magma freezes completely before much fractional crystallization has occurred, the magma becomes mafic igneous rock. But freezing of a magma that has been left over after lots of fractional crystallization has occurred produces felsic igneous rock.

Bowen’s Reaction Series

In the 1920s, the Canadian geologist Norman L. Bowen began a series of laboratory experiments designed to determine the sequence in which silicate minerals crystallize from a melt. First, Bowen melted powdered mafic igneous rock by raising its temperature to about 1280C. Then he cooled the melt just enough to cause part of it to solidify. Finally, he “quenched” the remaining melt by submerging it quickly in cold mercury. Quenching, which means sudden cooling to form a solid, transformed any remaining liquid into glass. The glass trapped the earlier-formed crystals within it. Bowen identified mineral crystals formed before quenching with a microscope, and he analyzed the chemical composition of the remaining glass. 

Bowen's reaction series indicates the succession of crystallization in cooling magma.

After experiments at different temperatures, Bowen found that, as new crystals form, they extract certain chemicals preferentially from the melt (figure above a). Thus, the chemical composition of the remaining melt progressively changes as the melt cools. Bowen described the specific sequence of mineral-producing reactions that take place in a cooling, initially mafic, magma. This sequence is now called Bowen’s reaction series in his honour.

Let’s examine the sequence more closely. In a cooling melt, olivine and calcium-rich plagioclase form first. This plagioclase reacts with the melt to form more, but different plagioclase; the plagioclase formed at a later stage contains more sodium (Na). Meanwhile, some olivine crystals react with the remaining melt to produce pyroxene, which may encase early olivine crystals or even replace them. However, some of the early olivine and Caplagioclase crystals settle out of the melt, taking iron, magnesium, and calcium atoms with them. By this process, the remaining melt becomes progressively enriched in silica. As the melt continues to cool, plagioclase continues to form, with later-formed plagioclase having progressively more sodium than earlier-formed plagioclase. Pyroxene crystals react with melt to form amphibole, and then amphibole reacts with the remaining melt to form biotite. All the while, crystals continue to settle out, so the remaining melt continues to become more felsic. At temperatures of between 650°C and 850°C, only about 10% melt remains, and this melt has a high silica content. At this stage, the final melt freezes, yielding quartz, K-feldspar (orthoclase), and muscovite. 

On the basis of his observations, Bowen realized that there are two tracks  to the reaction series. The “discontinuous” reaction series refers to the sequence  olivine, pyroxene, amphibole, biotite, K-feldspar-muscovite-quartz in that each step yields a different class of silicate mineral. The “continuous” reaction series refers to the progressive change from  calcium-rich to Na-rich plagioclase: the steps yield different versions of the  same mineral (figure above b). It’s important to note that not all minerals listed in the series appear in all igneous rock.  For example, a mafic magma may completely crystallize before felsic minerals such as quartz or K-feldspar have a chance to form.

Credits: Stephen Marshak (Essentials of Geology)

Studying Rock

Studying Rock

kacimo samy   2:04 م

Studying Rock 

Outcrop Observations 

The study of rocks begins by examining a rock in an outcrop. If the outcrop is big enough, such an examination will reveal relationships between the rock you’re interested in and the rocks around it, and will allow you to detect layering. Geologists carefully record observations about an outcrop, then break off a hand specimen, a fist-sized piece, that they can examine more closely with a hand lens (magnifying glass). Observation with a hand lens enables geologists to identify sand-sized or larger grains, and may enable them to describe the texture of the rock.

Thin-Section Study 

Studying rocks in thin section.

Geologists often must examine rock composition and texture in minute detail in order to identify a rock and develop a hypothesis for how it formed. To do this, they take a specimen back to the lab, make a very thin slice (about 0.03 mm thick, the thickness of a human hair) and mount it on a glass slide (figure above a–c). They study the resulting thin section with a petrographic microscope (petro comes from the Greek word for rock). A petrographic microscope differs from an ordinary microscope in that it illuminates the thin section with  transmitted polarized light. This means that the illuminating light beam first passes through a special polarity filter that makes all the light waves in the beam vibrate in the same plane, and then the light passes up through the thin section and then up through another polarizing filter. An observer looks through the thin section as if it were a window. When illuminated with transmitted polarized light, and viewed through two polarizing filters, each type of mineral grain displays a unique suite of colours (figure above d). The specific colour the observer sees depends on both the identity of the grain and its orientation with respect to the waves of polarized light. 

The brilliant colours and strange shapes in a thin section viewed in polarized light rival the beauty of an abstract painting or stained glass. By examining a thin section with a petrographic microscope, geologists can identify most of the minerals constituting the rock and can describe the way in which the grains connect to each other. They can make a record of the image by using a camera. A photograph taken through a petrographic microscope is called a photomicrograph.

High-Tech Analytical Equipment 

An electron micro-probe uses a beam of electrons to analyse the chemical composition of minerals.

Beginning in the 1950s, high-tech electronic instruments became available that enabled geologists to examine rocks on an even finer scale than is possible with a petrographic microscope. Modern research laboratories typically boast instruments such as electron micro-probes, which can focus a beam of electrons on a small part of a grain to create a signal that defines the  chemical  composition of the mineral (figure above); mass spectrometers, which analyse the proportions of atoms with different atomic weights contained in a rock; and X-ray diffractometers, which identify minerals by measuring how X-ray beams interact with crystals. Such instruments, in conjunction with optical examination, can provide geologists with highly detailed  characterizations of rocks, which in turn help them understand how the rocks formed and where the rocks came from. This information enables geologists to use the study of rocks as a basis for deciphering Earth history.

Credits: Stephen Marshak (Essentials of Geology)

The Basis of Rock Classification

The Basis of Rock Classification

kacimo samy   1:55 م

The Basis of Rock Classification 

Examples of three major rock groups.

Beginning in the 18th century, geologists struggled to develop a sensible way to classify rocks, for they realized, as did miners from centuries past, that not all rocks are the same. Classification schemes help us organize information and remember significant details about materials or objects, and they help us recognize similarities and differences among them. By the end of the 18th century, most geologists had accepted the genetic scheme for classifying rocks that we continue to use today. This scheme focuses on the origin (genesis) of rocks. Using this approach, geologists recognize three basic groups: (1) igneous rocks, which form by the freezing (solidification) of molten rock (figure above a); (2) sedimentary rocks, which form either by the cementing together of fragments (grains) broken off preexisting rocks or by the precipitation of mineral crystals out of water solutions at or near the Earth’s surface (figure above b); and (3) metamorphic rocks, which form when pre-existing rocks change character in response to a change in pressure and temperature conditions (figure above c). Metamorphic change occurs in the solid state, which means that it does not require melting. In the context of modern plate tectonics theory, different rock types form in different geologic settings (figure below).

A cross section illustrating various geologic settings in which rocks form.

Each of the three groups contains many different individual rock types, distinguished from one another by physical characteristics.

Describing grains in rock.

• Grain size: The dimensions of individual “grains” (here used in a general sense to mean fragments or crystals) in a rock may be measured in millimetres or centimetres. Some grains are so small that they can’t be seen without a microscope, whereas others are as big as a fist or larger. Some grains are equant, meaning that they have the same dimensions in all directions; some are inequant, meaning that the dimensions are not the same in all directions (figure above a, b). In some rocks, all the grains are the same size, whereas other rocks contain a variety of grain sizes.
• Composition: A rock is a mass of chemicals. The term rock composition refers to the proportions of different chemicals making up the rock. The proportion of chemicals, in turn, affects the proportion of different minerals  constituting the rock. 
•  Texture: This term refers to the arrangement of grains in a rock, that is, the way grains connect to one another and whether or not inequant grains are aligned parallel to each other. 
•  Layering: Some rock bodies appear to contain distinct layering, defined either by bands of different compositions or textures, or by the alignment of inequant grains so that they trend parallel to each other. Different types of  layering occur in different kinds of rocks. For example, the layering in sedimentary rocks is called bedding, whereas the layering in metamorphic rocks is called metamorphic foliation (figure below a, b). 

Layering a rock.

Each distinct rock type has a name. Names come from a variety of sources. Some come from the dominant component making up the rock, some from the region where the rock was first discovered or is particularly abundant, some from a root word of Latin origin, and some from a traditional name used by people in an area where the rock is found. 

Credits: Stephen Marshak (Essentials of Geology)

What Is Rock?

What Is Rock?

kacimo samy   7:59 ص

What Is Rock? 

To geologists, rock is a coherent, naturally occurring solid, consisting of an aggregate of minerals or, less commonly, of glass. Let’s take this definition apart to see what its components mean. 

• Coherent: A rock holds together, and thus must be broken to be separated into pieces. As a result of its coherence, rock can form cliff or can be carved into sculptures. A pile of unattached mineral grains does not constitute a rock. 
• Naturally occurring: Geologists consider only naturally occurring materials to be rocks, so manufactured materials, such as concrete and brick, do not qualify. 
• An aggregate of minerals or a mass of glass: The vast majority of rocks consist of an aggregate (a collection) of many mineral grains, and/or crystals, stuck or grown together. Some rocks contain only one kind of mineral, whereas others contain several different kinds. A few rock types consist of glass. 

Rocks, aggregates of mineral grains and/or crystals, can be clastic or crystalline.

What holds rock together? Grains in rock stick together to form a coherent mass either because they are bonded by  natural cement, mineral material that precipitates from water and fills the space between grains (figure above a), or because they i nterlock with one another like pieces in a jigsaw puzzle (figure above b). Rocks whose grains are stuck together by cement are called  clastic, whereas rocks whose crystals interlock with one another are called crystalline. Glassy rocks hold together because they originate as a continuous mass (that is, they have no separate grains), because glassy grains were welded together while still hot, or because they were cemented together at a later time. 

Types of rock exposures.

At the surface of the Earth, rock occurs either as broken chunks (pebbles, cobbles, or boulders) that have moved by falling down a slope or by being transported in ice, water, or wind, or as bedrock that is still attached to the Earth’s crust. Geologists refer to an exposure of bedrock as an outcrop. An outcrop may appear as a rounded knob out in a field, as a ledge forming a cliff or ridge, on the face of a stream cut (where running water dug down into bedrock), or along human-made roadcuts and excavations (figure above a–d). 

To people who live in cities or forests or on farmland, outcrops of bedrock may be unfamiliar, since bedrock may be completely covered by vegetation, sand, mud, gravel, soil, water, asphalt, concrete, or buildings. Outcrops are particularly rare in regions such as the midwestern United States, where, during the past million years, ice-age glaciers melted and buried  bedrock under thick deposits of debris. 

Credits: Stephen Marshak (Essentials of Geology)

Wegener’s Evidence for Continental Drift

Wegener’s Evidence for Continental Drift

kacimo samy   7:00 ص

Wegener’s Evidence for Continental Drift 

Wegener suggested that a vast supercontinent, Pangaea, existed until near the end of the Mesozoic Era (the interval of geologic time that lasted from 251 to 65 million years ago). He suggested that Pangaea then broke apart, and the landmasses moved away from each other to form the continents we see today. Let’s look at some of Wegener’s arguments and see what led him to formulate this hypothesis of continental drift.

The Fit of the Continents 

Almost as soon as maps of the Atlantic coastlines became available in the 1500s, scholars noticed the fit of the continents. The northwestern coast of Africa could tuck in against the eastern coast of North America, and the bulge of eastern South America could nestle cozily into the indentation of southwestern Africa. Australia, Antarctica, and India could all connect to the southeast of Africa, while Greenland, Europe, and Asia could pack against the northeastern margin of North America. In fact, all the continents could be joined, with remarkably few overlaps or gaps, to create Pangaea. Wegener concluded that the fit was too good to be coincidence and thus that the continents once did fit together.

Locations of Past Glaciations 

Glaciers are rivers or sheets of ice that flow across the land surface. As a glacier flows, it carries sediment grains of all sizes (clay, silt, sand, pebbles, and boulders). Grains protruding from the base of the moving ice carve scratches, called striations, into the substrate. When the ice melts, it leaves the sediment in a deposit called till, that buries striations. Thus, the occurrence of till and striations at a location serve as evidence that the region was covered by a glacier in the past (see chapter opening photo). By studying the age of glacial till deposits, geologists have determined that large areas of the land were covered by glaciers during time intervals of Earth history called ice ages. One of these ice ages occurred from about 326 to 267 Ma, near the end of the Paleozoic Era. 

Wegner's evidence for continental drift came from analysing the geologic record.

Wegener was an Arctic climate scientist by training, so it’s no surprise that he had a strong interest in glaciers. He knew that glaciers form mostly at high latitudes today. So he suspected that if he plotted a map of the locations of late Paleozoic glacial till and striations, he might gain insight into the locations of continents during the Paleozoic. When he plotted these locations, he found that glaciers of this time interval occurred in southern South America, southern Africa, southern India, Antarctica, and southern Australia. These places are now widely separated and, with the exception of Antarctica, do not currently lie in cold polar regions (figure above a). To Wegener’s amazement, all late Paleozoic glaciated areas lie adjacent to each other on his map of Pangaea. Furthermore, when he plotted the orientation of glacial striations, they all pointed roughly outward from a location in southeastern Africa. In other words, Wegener determined that the distribution of glaciations at the end of the Paleozoic Era could easily be explained if the continents had been united in Pangaea, with the southern part of Pangaea lying beneath the center of a huge ice cap. This distribution of glaciation could not be explained if the continents had always been in their present positions.

The Distribution of Climatic Belts 

If the southern part of Pangaea had straddled the South Pole at the end of the Paleozoic Era, then during this same time interval, southern North America, southern Europe, and northwestern Africa would have straddled the equator and would have had tropical or subtropical climates. Wegener searched for evidence that this was so by studying sedimentary rocks that were formed at this time, for the material making up these rocks can reveal clues to the past climate. For example, in the swamps and jungles of tropical regions, thick deposits of plant material accumulate, and when deeply buried, this material transforms into coal. And in the clear, shallow seas of tropical regions, large reefs develop. Finally, subtropical regions, on either side of the tropical belt, contain deserts, an environment in which sand dunes form and salt from evaporating seawater or salt lakes accumulates. Wegener speculated that the distribution of late Paleozoic coal, reef, sand-dune, and salt deposits could define climate belts on Pangaea. 

Sure enough, in the belt of Pangaea that Wegener expected to be equatorial, late Paleozoic sedimentary rock layers include abundant coal and the relicts of reefs. And in the portions of Pangaea that Wegener predicted would be subtropical, late Paleozoic sedimentary rock layers include relicts of desert dunes and deposits of salt (figure above b). On a present-day map of our planet, exposures of these ancient rock layers scatter around the globe at a variety of latitudes. On Wegener’s Pangaea, the exposures align in continuous bands that occupy appropriate latitudes.

The Distribution of Fossils 

Today, different continents provide homes for different species. Kangaroos, for example, live only in Australia. Similarly, many kinds of plants grow only on one continent and not on others. Why? Because land-dwelling species of animals and plants cannot swim across vast oceans, and thus evolved independently on different continents. During a period of Earth history when all continents were in  contact, however, land animals and plants could have migrated among many continents. 

With this concept in mind, Wegener plotted fossil occurrences of land-dwelling species that existed during the late Paleozoic and early Mesozoic Eras (between about 300 and 210 million years ago) and found that these species had indeed existed on several continents (figure above c). Wegener argued that the distribution of fossil species required the continents to have been adjacent to one another in the late Paleozoic and early Mesozoic Eras.

Matching Geologic Units 

 Further evidence of drift: rocks on different sides of the ocean match.

An art historian can recognize a Picasso painting, an architect knows what makes a building look “Victorian,” and a geoscientist can identify a distinctive assemblage of rocks. Wegener found that the same distinctive Precambrian rock assemblages occurred on the eastern coast of South America and the western coast of Africa, regions now separated by an ocean  (figure above a). If the continents had been joined to create Pangaea in the past, then these matching rock groups would have been adjacent to each other, and thus could have composed continuous blocks or belts. Wegener also noted that features of the Appalachian Mountains of the United States and Canada closely resemble mountain belts in southern Greenland, Great Britain, Scandinavia, and northwestern Africa (figure above b, c), regions that would have lain adjacent to each other in Pangaea. Wegener thus demonstrated that not only did the coastlines of continents match, so too did the rocks adjacent to the coastlines.

Criticism of Wegener’s Ideas 

Wegener’s model of a supercontinent that later broke apart explained the distribution of ancient glaciers, coal, sand dunes, rock assemblages, and fossils. Clearly, he had compiled a strong circumstantial case for continental drift. But as noted earlier, he could not adequately explain how or why continents drifted. He left on his final expedition to Greenland having failed to convince his peers, and he died without knowing that his ideas, after lying dormant for decades, would be reborn as the basis of the broader theory of plate tectonics. 

In effect, Wegener was ahead of his time. It would take three more decades of research before geologists obtained sufficient data to test his hypotheses properly. Collecting this data required instruments and techniques that did not exist in Wegener’s day. Of the many geologic discoveries that ultimately opened the door to plate tectonics, perhaps the most important came from the discovery of a phenomenon called paleomagnetism, so we discuss it next.

Credit: Stephen Marshak (Essentials of Geology)

What Are Earth Layers Made Of?

What Are Earth Layers Made Of?

kacimo samy   7:07 ص

What Are Earth Layers Made Of? 

A modern view of Earth‘s interior layers.

As a result of studies during the past century, geologists have a pretty clear sense of what the layers inside the Earth are made of. Let’s now look at the properties of individual layers in more detail (figure above a, b).

The Crust 

When you stand on the surface of the Earth, you are standing on top of its outermost layer, the crust. The crust is our home and the source of all our resources. How thick is this all important layer? Or, in other words, what is the depth to the crust-mantle boundary? An answer came from the studies of Andrija Mohorovicˇic´, a researcher working in Zagreb, Croatia. In 1909, he discovered that the velocity of earthquake waves suddenly increased at a depth of tens of kilometres beneath the Earth’s surface, and he suggested that this increase was caused by an abrupt change in the properties of rock. Later studies showed that this change can be found most everywhere around our planet, though it occurs at different depths in different locations. Specifically, it’s deeper beneath continents than beneath oceans. Geologists now consider the change to define the base of the crust, and they refer to it as the Moho in Mohorovicˇic´’s honour. The relatively shallow depth of the Moho (7 to 70 km, depending on location) as compared to the radius of the Earth (6,371 km) emphasizes that the crust is very thin indeed. In fact, the crust is only about 0.1% to 1.0% of the Earth’s radius, so if the Earth were the size of a balloon, the crust would be about the thickness of the balloon’s skin.

The crust is not simply cooled mantle, like the skin on chocolate pudding, but rather consists of a variety of rocks that differ in composition (chemical make-up) from mantle rock. Geologists distinguish between two fundamentally different types of crust oceanic crust, which underlies the sea floor, and continental crust, which underlies continents. 

Oceanic crust is only 7 to 10 km thick. At highway speeds (100 km per hour), you could drive a distance equal to the thickness of the oceanic crust in about five minutes. At the top, we find a blanket of sediment, generally less than 1 km thick, composed of clay and tiny shells that settled like snow out of the sea. Beneath this blanket, the oceanic crust consists of a layer of basalt and, below that, a layer of gabbro. 

 A table and a graph illustrating the abundance of elements in the Earth’s crust.

Most continental crust is about 35 to 40 km thick about four to five times the thickness of oceanic crust but its thickness varies significantly. In some places, continental crust has been stretched and thinned so it’s only 25 km from the surface to the Moho, and in some places, the crust has been crumpled and thickened to become up to 70 km thick. In contrast to oceanic crust, continental crust contains a  great variety of rock types, ranging from mafic to felsic in composition. On average, upper continental crust is less mafic than oceanic crust it has a felsic (granite-like) to intermediate composition so continental crust overall is less dense than oceanic crust. Notably, oxygen is the most abundant  element in the crust (figure above).

The Mantle 

The mantle of the Earth forms a 2,885-km-thick layer surrounding the core. In terms of volume, it is the largest part of the Earth. In contrast to the crust, the mantle consists entirely of an ultramafic (dark and dense) rock called peridotite. This means that peridotite, though rare at the Earth’s surface, is actually the most abundant rock in our planet! Researchers have found that earthquake-wave velocity changes at a depth of 400 km and again at a depth of 660 km in the mantle. Based on this observation, they divide the mantle into two sublayers: the upper mantle, down to a depth of 660 km, and the lower mantle, from 660 km down to 2,900 km. The transition zone is the interval between 400 km and 660 km deep. 

Almost all of the mantle is solid rock. But even though it’s solid, mantle rock below a depth of 100 to 150 km is so hot that it’s soft enough to flow. This flow, however, takes place extremely slowly at a rate of less than 15  cm a year. Soft here does not mean liquid; it simply means that over long periods of time mantle rock can change shape without breaking. We stated earlier that almost all of the mantle is solid. We used the word “almost” because up to a few percent of the mantle has melted. This melt occurs in films or bubbles between grains in the mantle at a depth of 100 to 200 km beneath the ocean floor. Although overall, the temperature of the mantle increases with depth, temperature also varies significantly with location even at the same depth. The warmer regions are less dense, while the cooler regions are denser. The distribution of warmer and cooler mantle indicates that the mantle convects like water in a simmering pot; warmer mantle is relatively buoyant and gradually flows upward, while cooler, denser mantle sinks.

The Core 

Early calculations suggested that the core had the same density as gold, so for many years people held the fanciful hope that vast riches lay at the heart of our planet. Alas, geologists eventually concluded that the core consists of a far less glamorous material, iron alloy (iron mixed with tiny amounts of other elements). Studies of seismic waves led geo scientists to divide the core into two parts, the outer core (between 2,900 and 5,155 km deep) and the inner core (from a depth of 5,155 km down to the Earth’s centre at 6,371 km). The outer core consists of liquid iron alloy. It can exist as a liquid because the temperature in the outer core is so high that even the great pressures squeezing the region cannot keep atoms locked into a solid framework. The iron alloy of the outer core can flow, and this flow generates Earth’s magnetic field. 

The inner core, with a radius of about 1,220 km, is a solid iron alloy that may reach a temperature of over 4,700°C. Even though it is hotter than the outer core, the inner core is a solid because it is deeper and is subjected to even greater pressure. The pressure keeps atoms locked together tightly in very dense materials.

The Lithosphere and the Asthenosphere 

So far, we have identified three major layers (crust, mantle, and core) inside the Earth that differ compositionally from each other. Earthquake waves travel at different velocities through these layers. An alternative way of thinking about Earth layers comes from studying the degree to which the material making up a layer can flow. In this context, we distinguish between rigid materials, which can bend or break but cannot flow, and plastic materials, which are relatively soft and can flow without breaking.

A block diagram of the lithosphere, emphasizing the difference between continental and oceanic lithosphere.

Geologists have determined that the outer 100 to 150  km of the Earth is relatively rigid. In other words, the Earth has an outer shell composed of rock that cannot flow easily. This outer layer is called the lithosphere, and it consists of the crust plus the uppermost, cooler part of the mantle. We refer to the portion of the mantle within the lithosphere as the lithospheric mantle. Note that the terms lithosphere and crust are not synonymous the crust is just the upper part of the lithosphere. The lithosphere lies on top of the asthenosphere, which is the portion of the mantle in which rock can flow. The boundary between the lithosphere and asthenosphere occurs where the temperature reaches about 1280°C, for at temperatures higher than this value mantle rock becomes soft enough to flow. 

Geologists distinguish between two types of lithosphere (figure above). Oceanic lithosphere, topped by oceanic crust, generally has a thickness of about 100 km. In contrast, continental lithosphere, topped by continental crust, generally has a thickness of about 150 km. Notice that the asthenosphere is entirely in the mantle and generally lies below a depth of 100 to 150 km. We can’t assign a specific depth to the base of the asthenosphere because all of the mantle below 150 km can flow, but for convenience, some geologists consider the base of the asthenosphere to be the top of the transition zone.
Credits: Stephen Marshak (Essentials of geology)