Carbonate Petrography

Carbonate petrography is the study of limestones, dolomites and associated deposits under optical or electron microscopes greatly enhances field studies or core observations and can provide a frame of reference for geochemical studies.

25 strangest Geologic Formations on Earth

The strangest formations on Earth.

What causes Earthquake?

Of these various reasons, faulting related to plate movements is by far the most significant. In other words, most earthquakes are due to slip on faults.

The Geologic Column

As stated earlier, no one locality on Earth provides a complete record of our planet’s history, because stratigraphic columns can contain unconformities. But by correlating rocks from locality to locality at millions of places around the world, geologists have pieced together a composite stratigraphic column, called the geologic column, that represents the entirety of Earth history.

Folds and Foliations

Geometry of Folds Imagine a carpet lying flat on the floor. Push on one end of the carpet, and it will wrinkle or contort into a series of wavelike curves. Stresses developed during mountain building can similarly warp or bend bedding and foliation (or other planar features) in rock. The result a curve in the shape of a rock layer is called a fold.

Friday, 27 November 2015

The Proterozoic: The Earth in Transition

Growth of Continents 

The Proterozoic Eon spans roughly 2 billion years, from about 2.5 Ga to the beginning of the Cambrian Period at 542 Ma thus, it encompasses almost half of Earth’s history. During Proterozoic time, Earth’s surface environment changed from being an unfamiliar world of fast-moving plates, small continents, and an oxygen-poor atmosphere, to the more familiar world of slower plates, large continents, and an oxygen-rich atmosphere. First, let’s look at changes to the continents. New continental crust continued to form during the Proterozoic Eon, but at progressively slower rates in fact, by the middle of the eon, over 90% of the Earth’s continental crust had formed. As Archean proto-continents collided with each other and with volcanic island arcs and hot-spot volcanoes, still larger continents gradually assembled. Significantly, the interiors of these larger continents became isolated from heating by subduction-related igneous activity that happened along its margins. Interior regions, therefore, slowly cool and strengthen until they become very rigid and durable. Such a region of cold, relatively stable continental crust is called a craton. All cratons that exist today had formed by about 1 Ga (figure above); therefore, crust in cratons (the old parts of continents) ranges from about 3.85 Ga to about 1 Ga. 

To understand the character of a craton, let’s examine North America’s craton a bit more closely. We see that it consists of two regions (figure above). Throughout the shield, outcrops expose Precambrian “basement,” which consists of igneous and metamorphic rocks older than about 1 Ga. The landscape of the shield tends to have fairly low relief there are small hills and valleys, but no dramatic mountain ranges. Most of North America’s shield lies in Canada, so geologists refer to it as the Canadian Shield. Throughout the cratonic platform, which surrounds the shield and also underlies Hudson Bay, a blanket of Paleozoic or Mesozoic cover strata overlies the Precambrian basement. 

By using isotopic dating on samples from both outcrops and drill holes, geologists have been able to subdivide the basement of North America’s craton into distinct blocks or provinces, each of which has been given a name (figure above). It appears that the Canadian Shield consists of several Archean crust blocks sutured together by Proterozoic orogens. The basement of the cratonic platform in the United States, in contrast, grew when a series of volcanic island arcs and continental slivers “accreted” (attached) to the margin of the Canadian Shield between  1.8 and 1.6 Ga. In the Midwest, granite plutons intruded much of this accreted region, and rhyolite ash flows covered it, between 1.5 and 1.3 Ga. Successive collisions ultimately brought together most continental crust on Earth into a single supercontinent, named Rodinia, by around 1 Ga. The last major collision during the formation of Rodinia produced a large orogen called the Grenville orogen. 

If you look at a popular (though not universally accepted) reconstruction of Rodinia, you can identify the crustal provinces that would eventually become the familiar continents of today (figure above a). Several studies suggest that sometime between 800 and 600 Ma, Rodinia “turned inside out,” in that Antarctica, India, and Australia broke away from western North America and swung around and collided with the future South America, possibly forming a short-lived supercontinent that some geologists refer to as Pannotia (figure above b). The map of the Earth clearly changed radically during the Proterozoic. But that’s not all that changed fossil evidence suggests that this eon also saw important steps in the evolution of life. When the Proterozoic began, most life was prokaryotic, meaning that it consisted of single-celled organisms (archaea and bacteria) without a nucleus. Though studies of chemical fossils hint that eukaryotic life, consisting of cells that have nuclei, originated as early as 2.7 Ga, the first possible body fossil of a eukaryotic organism occurs in 2.1-Ga rocks, and abundant body fossils of eukaryotic organisms can be found only in rocks younger than about 1.2 Ga. Thus the proliferation of eukaryotic life, the foundation from which complex organisms eventually evolved, took place during the Proterozoic. The last half-billion years of the Proterozoic Eon saw the remarkable transition from simple organisms into complex ones. Ciliate protozoans (single-celled organisms coated with fibers that give them mobility) appear at about 750 Ma. 

A great leap forward in complexity of organisms occurred during the next 150 million years of the eon, for sediments deposited perhaps as early as 620 Ma and certainly by 565 Ma contain several types of multicellular organisms that together constitute the Ediacaran fauna, named for a region in southern Australia where fossils of these organisms were first found. Ediacaran species survived into the beginning of the Cambrian before becoming extinct. Their fossil forms suggest that some of these invertebrate organisms resembled jellyfish, while others resembled worms (figure above a). The evolution of life played a key role in the evolution of Earth’s atmosphere. Before life appeared, there was hardly any free oxygen (O2) in the atmosphere. With the appearance of photosynthetic organisms, oxygen began to enter the  atmosphere. But it was not until about 2.4 Ga that the concentration of oxygen in the atmosphere increased dramatically. This event, called the great oxygenation event, happened when other environments were no longer able to absorb or dissolve the oxygen produced by organisms, so the oxygen began to accumulate as a gas in air. As a result, the oceans became oxidizing environments that, for reasons described in chemistry books, could no longer contain large quantities of dissolved iron. Between 2.4 Ga and 1.8 Ga, huge amounts of iron settled out of the ocean to form colorful sedimentary beds known as banded iron formation (BIF). BIF consists of alternating layers of iron oxide minerals (hematite or magnetite) and jasper (red chert) as illustrated by the chapter opening photo. Radical climate shifts occurred on Earth at the end of the Proterozoic Eon. Specifically, accumulations of glacial sediments occur worldwide in late Proterozoic stratigraphic sequences. What’s strange about the occurrence of these sediments is that they can be found even in regions that were located at the equator during these times (figure above b). This observation implies that the entire planet was cold enough for glaciers to form at the end of the Proterozoic. Geologists still are debating the history of these global ice ages; in one model, glaciers covered the land and perhaps the entire ocean surface froze, resulting in snowball Earth (figure above c). The shell of ice cut off the oceans from the atmosphere, causing oxygen levels in the sea to drop drastically, so many life forms died off. What brought an end to snowball Earth conditions? According to one model, the icy sheath also prevented atmospheric CO2 from dissolving in seawater, but it did not prevent volcanic activity from continuing to add CO2 to the atmosphere. Earth would have remained a snowball forever, were it not for the addition of volcanic CO2, a greenhouse gas that traps heat in the atmosphere much as glass panes trap heat in a greenhouse. As the CO2 concentration increased, Earth warmed up and eventually the glaciers melted. 

Introducing the Phanerozoic Eon 

As the Proterozoic came to a close, Earth’s climate warmed and continents drifted apart life evolved and diversified to occupy the new environments that formed. Over a relatively short period of time, shells appeared and the fossil record became much more complete. This event defines the end of the Proterozoic Eon and therefore, of the Precambrian, and the start of the Phanerozoic Eon. Of note, geologists recognized the siginificance of this event long before they could assign it a numerical age (currently 542 Ma). The Phanerozoic Eon consists of three eras the Paleozoic (Greek for ancient life), the Mesozoic (middle life), and the Cenozoic (recent life). Geologists have divided the Mesozoic and Cenozoic each into three periods and the Paleozoic into six periods. In the sections that follow, we consider changes in the map of our planet’s surface (its paleogeography), as manifested by the distribution of continents, seas, and mountain belts, as well as life evolution that happened during the three eras.
Credits: Stephen Marshak (Essentials of Geology)

The Archean Eon: Birth of the Continents and Life

The Archean Eon

The Archean Eon has given us where we lives now, land.

Land Appears 

The boundary between the Archean (from the Greek word for beginning) Eon and the Hadean Eon occurs at about 3.85 Ga. Effectively, this date marks the time at which substantial quantities of crustal rocks, including rocks that originated as marine sediments, formed. With the advent of the Archean, crust was locally cool and stable enough for rocks to survive and for isotopic clocks to start ticking. Geologists still debate about whether plate tectonics in the form that occurs today operated in the early part of the Archean Eon. Most researchers picture an early Archean Earth with rapidly moving small plates, numerous volcanic island arcs, and abundant hot-spot volcanoes. Others propose that early Archean lithosphere was too warm and buoyant to subduct, and that plate tectonics could not have operated until the later part of the Archean or later; these authors argue that plume-related volcanism or some other process was the main source of new crust until the late Archean. Regardless of which model ultimately proves more accurate, it is clear that the Archean was a time during which significant volumes of new continental crust were generated. What processes produced continental crust? According to one model, early crust formed from mafic igneous rocks that originally extruded or intruded at convergent plate boundaries and/or hot-spot volcanoes. Once formed, these rocks were too buoyant to be subducted, so when the arcs and plateaus collided with one another, they sutured together to form larger blocks that remained at the Earth’s surface. The development of convergent plate boundaries along the margins of these blocks, and of rifts and hot spots within the blocks, led to production of flood basalts. Partial melting of basaltic crust yielded felsic and intermediate rocks. As collisions continued, the blocks coalesced into still larger proto-continents (figure above a, b), which slowly cooled and became stronger. As a result of these processes, the first long-lived blocks of durable continental crust came into existence between 3.2 and 2.7 Ga, and by the end of the Archean Eon, about 80% of the Earth’s continental area had formed (figure above c). A clear stratigraphic record of marine sediment deposition appears in remnants of Archean crust, indicating that oceans filled in the Archean and have existed ever since. Presumably, permanent oceans could survive only after the Earth’s surface had cooled below the boiling point of water. Prior to that time, gaseous H2O saturated the atmosphere in fact, prior to ocean formation, H2O and CO2 were the dominant gases of the atmosphere. Once the oceans formed, however, the atmosphere lost most of its H2O. And once liquid water existed, most atmospheric CO2 dissolved into it, so CO2 also went from being a major component of the atmosphere to being a trace component. Thus, the Archean saw the atmosphere change from being a foggy mixture of H2O and CO2 into being a transparent gas dominantly composed of nitrogen (N2) gas. Since nitrogen is inert (it doesn't chemically react with or dissolve in other materials), it was left behind.

The First Life 

Clearly, the Archean Eon saw many firsts in Earth history. Not only did the first continents appear during the Archean, but probably also the first life. The search for the earliest evidence of life continues to make headlines in the popular media. Most geologists currently conclude that life has existed on Earth since at least  3.5 Ga, and perhaps since 3.8 Ga, for rocks of this age contain chemical signatures of organisms. The oldest undisputed fossil forms of bacteria and archaea occur in 3.2-Ga rocks  (figure above a). (Shapes resembling such organisms occur in rocks as old as 3.5 Ga, but their identity remains less certain.) Archean strata at some localities contain stromatolites, distinctive layered mounds of sediment. Stromatolites that developed after about 3.2 Ga form because cyanobacteria secrete a mucuslike substance to which sediment settling from water sticks. As the mat gets buried, new cyanobacteria colonize the top of the sediment, building a mound upward (figure above b); modern examples locally occur in shallow, tropical waters. What specific environment on the Archean Earth served as the cradle of life? Laboratory experiments conducted in the 1950s led many researchers to think that life began in warm pools of surface water, beneath a methane- and ammoniarich atmosphere streaked by bolts of lightning. More recent researchers suggest instead that submarine hot-water vents, so-called black smokers, served as the hosts of the first organisms. These vents emit clouds of ion-charged solutions from which sulphide minerals precipitate and build chimneys. The earliest life in the Archean Eon may well have been thermophilic (heat-loving) bacteria or archaea that dined on pyrite at dark depths in the ocean alongside these vents. Later in the Archean, organisms evolved the ability to carry out photosynthesis, and moved into shallower, well lit water. As the Archean Eon came to a close, the first continents had formed, and life had colonized not only the depths of the sea but also the shallow marine realm. Plate tectonics had commenced, continental drift was taking place, collisional mountain belts were forming, and erosion was occurring. The atmosphere was gradually accumulating oxygen, although probably this gas still accounted for only a very small percentage of the air; Archean air was unbreathable. The stage was set for another major change in the Earth System.
Credits: Stephen Marshak (Essentials of Geology)

The Hadean Eon: Before the Rock Record

Hadean Eon

James Hutton, the 18th-century Scottish geologist who was the first to provide convincing evidence that the Earth was very old, could not measure Earth’s age directly, and indeed speculated that there may be “no vestige of a beginning.” But isotopic dating studies conducted in recent decades have shown that it is possible, in fact, to assign a numerical age to our planet’s formation. Specifically, dates obtained for a class of meteorites thought to be remnants of the planetesimal cloud out of which the Earth formed yield an age of 4.57 Ga. Geologists currently take this age to be the Earth’s birth date. But the oldest whole rock yet found is only 4.03 Ga, and a clear record of Earth history, as recorded in continental crustal rocks, does not begin until about 3.85 Ga. Geologists refer to the mysterious time interval between the birth of Earth and 3.85 Ga as the Hadean Eon (from Hades, the Greek god of the underworld) because during this interval the Earth’s surface was, at times, like an inferno. Many major events happened during the Hadean. By about 4.5 Ga, the Earth underwent internal  differentiation gravity pulled molten iron down to the center of the Earth, where it accumulated to form the core, leaving a mantle composed of ultramafic rock. Researchers suggest that soon after or perhaps during differentiation, a Mars-sized protoplanet collided with the Earth. The energy of this collision blasted away a significant fraction of Earth’s mantle; the resulting debris formed a ring of silicate-rock debris orbiting the Earth. This ring then coalesced to form the Moon, which, when first formed, was less than 20,000 km away. (By comparison, the Moon is 384,000 km from Earth today.)

This speculative painting depicts the Hadean Earth's surface as a magma ocean pummeled by meteorites. The Moon was much closer then, but might not have been visible through the dense atmosphere.
In the wake of differentiation and Moon formation, the Earth was so hot that much of its surface was an ocean of seething magma (figure above). Rafts of solid rock formed temporarily on the surface of the magma ocean, but these eventually sank and remelted. This stage lasted at least until about 4.4 Ga. After that time, the amount of radioactive heat generation decreased (because elements with short half-lives had decayed), so the Earth might have become cool enough for solid rocks to form at its surface. The evidence for this statement comes from western Australia, where geologists have found 4.4-Ga grains of a durable mineral called zircon. During the Hadean Eon, outgassing of the Earth’s mantle began to take place. This means that volatile (gassy) elements or compounds originally incorporated in mantle minerals were released and bubbled out of volcanoes, along with lava. The gases accumulated into a toxic atmosphere of water (H2O), methane (CH4), ammonia (NH3), hydrogen (H2), nitrogen (N2), carbon dioxide (CO2), sulfur dioxide (SO2), and other gases. Some researchers speculate that gases from comets colliding with Earth may have contributed additional gases to the early atmosphere. If the Hadean Earth’s surface was sufficiently cool for an extensive solid crustal rock to form beginning at 4.4 Ga, then the first oceans may have accumulated soon thereafter, when water in the atmosphere condensed and fell as rain. Though mineral grains as old as 4.4 Ga exist, the oldest whole rock yet found on Earth has an age of only 4.03 Ga, and most rocks are younger than 3.85 Ga. What destroyed most, if not all, of the pre-3.85-Ga rock (and oceans, if they existed) of the Earth? The answer comes from studies of cratering on the Moon. These studies suggest that the Moon and, therefore, all inner planets of the Solar System underwent intense meteor bombardment between 4.0 and 3.85 Ga. Researchers speculate that this bombardment would have pulverized and/ or melted almost all crust that had existed on Earth at the time, and would have destroyed the existing atmosphere and ocean. Only after the bombardment ceased could long-lasting crust, atmosphere, and oceans begin to form. The discovery of 3.85-Ga marine sedimentary rocks in Greenland suggests that the appearance of land and sea happened quite soon after bombardment ceased. What did the Earth’s surface look like at 3.85 Ga? An observer probably would have found small, barren landmasses, spotted with volcanoes, poking up above an acidic sea. But both land and sea would have been obscured by murky, dense (CO2- and SO2-rich) air.
Credits: Stephen Marshak (Essentials of Geology)

Numerical age and geologic time

Dating Sedimentary Rocks? 

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

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

The Geologic Time Scale 

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

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

What Is the Age of the Earth? 

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

Picturing Geologic Time 

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

Wednesday, 25 November 2015

How do we determine numerical age of Earth?

Numerical age determination

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

Radioactive Decay 

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

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

Isotopic Dating 

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

Table above lists particularly useful elements. Each radioactive element has its own half-life. (Note that carbon dating is not used for dating rocks because appropriate carbon isotopes occur only in organisms and radioactive carbon has a very short half-life). Second, we must identify the right kind of minerals to work with. Not all minerals contain radioactive elements, but fortunately some fairly common minerals do. Once we have found the right kind of minerals, we can set to work using the following steps. 
  • Collecting the rocks: We need to find un-weathered rocks for dating, for the chemical reactions that happen during weathering may lead to the loss of some isotopes. 
  • Separating the minerals: The rocks are crushed, and the appropriate minerals are separated from the debris. 
  • Extracting parent and daughter isotopes: To separate out the parent and daughter isotopes from minerals, we can use several techniques, including dissolving the minerals in acid or evaporating portions of them with a laser. 
  • Analyzing the parent-daughter ratio: Once we have a sample of appropriate atoms, we pass them through a mass spectrometer, an instrument that uses a strong magnet to separate isotopes from one another according to their respective weights (figure below). The instrument can count the number of atoms of specific isotopes separately. 

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

What Does an Isotopic Date Mean? 

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

Other Methods of Determining Numerical Age

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

Sunday, 22 November 2015

How do Earthquakes causes damage?

Damages from Earthquakes

An area ravaged by a major earthquake is a heartbreaking sight. The terror and sorrow etched on the faces of survivors mirror the inconceivable destruction. This destruction comes as a result of many processes.

Ground Shaking and Displacement 

An earthquake starts suddenly and may last from a few seconds to a few minutes. Different kinds of earthquake waves cause different kinds of ground motion (figure above). The nature and severity of the shaking at a given location depend on four factors: 
  1. the magnitude of the earthquake, because larger magnitude events release more energy; 
  2. the distance from the focus, because earthquake energy decreases as waves pass through the Earth; 
  3. the nature of the substrate at the location (that is, the character and thickness of different materials beneath the ground surface) because earthquake waves tend to be amplified in weaker substrate; and 
  4. the “frequency” of the earthquake waves (where frequency equals the number of waves that pass a point in a specified interval of time). 

If you’re out in an open field during an earthquake, ground motion alone won’t kill you, you may be knocked off your feet and bounced around a bit, but your body is too flexible to break. Buildings and bridges aren't so lucky (figure above a-d). When earthquake waves pass, they sway, twist back and forth, or lurch up and down, depending on the type of wave motion. As a result, connectors between the frame and facade of a building may separate, so the facade crashes to the ground. The flexing of walls shatters windows and makes roofs collapse. Floors or bridge decks may rise up and slam down on the columns that support them, thereby crushing the columns. Some buildings collapse with their floors piling on top of one another like pancakes in a stack, some crumble into fragments, and some simply tip over. The majority of  earthquake-related deaths and injuries happen when people are hit by debris or are crushed beneath falling walls or roofs. Aftershocks worsen the problem, because they may topple already weakened buildings, trapping rescuers. During earthquakes, roads, rail lines, and pipelines may also buckle and rupture. If a building, fence, road, pipeline, or rail line straddles a fault, slip on the fault can crack the structure and separate it into two pieces.


The shaking of an earthquake can cause ground on steep slopes or ground underlain by weak sediment to give way. This movement results in a landslide, the tumbling and flow of soil and rock down-slope. Earthquake triggered landslides occur commonly along the coast of  California where expensive homes perch on steep cliffs looking out over the Pacific. When the cliffs collapse, the homes may tumble to the beach below (figure above a and b). Such events lead to the misperception that “California will someday fall into the sea.” Although small portions of the coastline do collapse, the state as a whole remains firmly attached to the continent, despite what  Hollywood  scriptwriters say.

Sediment Liquefaction 

In 1964, an MW 7.5 earthquake struck Niigata, Japan. A portion of the city had been built on land underlain by wet sand. During the ground shaking, foundations of over 15,000 buildings sank into their substrate, causing walls and roofs to crack. Several four-story-high buildings in a newly built apartment complex tipped over (figure above c). The same year, on Good Friday, an MW  9.2 earthquake devastated southern Alaska. In the Turnagain Heights neighbourhood of Anchorage, the event led to catastrophe. The neighbourhood was built on a small terrace of uplifted sediment. The edge of the terrace was a 20-m high escarpment that dropped down to Cook Inlet, a bay of the Pacific Ocean. As the ground shaking began, a layer of wet clay beneath the development turned into mud, and when this happened, the overlying layers of sediment, along with the houses built on top of them, slid seaward. In the process, the layers broke into separate blocks that tilted, turning the landscape into a chaotic jumble, and resulting in the destruction of the neighbourhood (figure above d). 

In 2011, an earthquake in Christchurch, New Zealand, caused sand to erupt and produce small, cone-shaped mounds on the ground surface  (figure above a). The transfer of sand from underground onto the surface led to formation of depressions large enough to  swallow cars (figure above b). All of these examples are manifestations of a phenomenon called sediment liquefaction. During liquefaction, pressure in the water filling the pores between grains in wet sand push the grains apart so that they become surrounded by water and no longer rest against each other. In wet clay, shaking breaks the weak electrostatic charges that hold clay flakes together, so what had been a gel-like, stable mass becomes slippery mud. As the material above the liquefied sediment settles downward, pressure can squeeze the sand upward and out onto the ground surface. The resulting cone-shaped mounds are variously known as sand volcanoes, sand boils, or sand blows. The settling of sedimentary layers down into a liquefied layer can also disrupt bedding and can lead to formation of open fissures of the land surface (figure above c).


The shaking during an earthquake can make lamps, stoves, or candles with open flames tip over, and it may break wires or topple power lines, generating sparks. As a consequence, areas already turned to rubble, and even areas not so badly damaged may be consumed by fire. Ruptured gas pipelines and oil tanks feed the flames, sending columns of fire erupting skyward (figure above). Fire fighters might not even be able to reach the fires, because the doors to the fire house won’t open or rubble blocks the streets. Moreover, fire fighters may find themselves without water, for ground shaking and landslides damage water lines. Once a fire starts to spread, it can become an unstoppable inferno. Most of the destruction of the 1906 San Francisco earthquake, in fact, resulted from fire. For three days, the blaze spread through the city until fire fighters contained it by blasting a fire break. By then, 500 blocks of structures had turned to ash, causing 20 times as much financial loss as the shaking itself. When a large earthquake hit Tokyo in 1923, fires set by cooking stoves spread quickly through the wood-and-paper buildings, creating an inferno a “fire storm” that heated the air above the city. As hot air rose, cool air rushed in, creating wind gusts of over 100 mph, which stoked the blaze and incinerated 120,000 people.


The azure waters and palm fringed islands of the Indian Ocean’s east coast hide one of the most seismically active plate boundaries on Earth, the Sunda Trench. Along this convergent boundary, the Indian Ocean floor subducts at about 4 cm per year, leading to slip on large thrust faults. Just before 8:00 a.m. on December 26, 2004, the crust above a 1,300-km long by 100-km-wide portion of one of these faults lurched westward by as much as 15 m. The break started at the hypocentre and then propagated north at 2.8 km/s; thus, the rupturing took 9 minutes. This slip triggered a great earthquake (MW 9.3) and pushed the sea floor up by tens of centimetres. The rise of the sea floor, in turn, shoved up the overlying water. Because the area that rose was so broad, the volume of displaced water was immense. As a consequence, a tragedy of an unimaginable extent was about to unfold. Water from above the up-thrust sea floor began moving outward from above the fault zone, a process that generated a series of giant waves travelling at speeds of about 800 km per hour (500  mph) almost the speed of a jet plane (figure above a). Geologists now use the term tsunami for a wave produced by displacement of the sea floor. The displacement can be due to an earthquake, submarine landslide, or volcanic explosion. Tsunami is a Japanese word that translates literally as harbour wave, an apt name because tsunamis can be particularly damaging to harbour towns. In older literature such waves were called “tidal waves,” because when one arrives, water rises as if a tide were coming in, but in fact the waves have nothing to do with daily tidal cycles. Regardless of cause, tsunamis are very different from familiar, wind-driven storm waves. Large wind-driven waves can reach heights of 10 to 30 meters in the open ocean. But even such monsters have wavelengths of only tens of meters, and thus contain a relatively small volume of water. In contrast, although a tsunami in deep water may cause a rise in sea level of at most only a few tens of centimetres a ship crossing one wouldn't even notice tsunamis have wavelengths of tens to hundreds of kilometres and an individual wave can be several kilometres wide, as measured perpendicular to the wave front. Thus, the wave involves a huge volume of water. In simpler terms, we can think of the width of a tsunami, in map view, as being more than 100 times the width of a wind-driven wave. Because of this difference, a storm wave and a tsunami have very different effects when they strike the shore. When a wave approaches the shore, friction between the base of the wave and the sea floor slows the bottom of the wave, so the back of the wave catches up to the front, and the added volume of water builds the wave higher (figure above b). The top of the wave may fall over the front of the wave and cause a breaker. In the case of a wind-driven wave, the breaker may be tall when it washes onto the beach, but because the wave doesn't contain much water, the wave runs out of water and friction slows it to a stop on the beach. Then, gravity causes the water to spill back seaward. In the case of a tsunami, the wave is so wide that, as friction slows the wave, it builds into a “plateau” of water that can be tens of meters high, many kilometres wide, and hundreds of kilometres long. Thus, when a tsunami reaches shore, it contains so much water that it crosses the beach and, if the land is low-lying, just keeps on going, eventually covering a huge area (figure above c). 

Tsunami damage can be catastrophic. The December 2004 waves struck Banda Aceh, a city at the north end of the island of Sumatra, on a beautiful, cloudless day (figure above a). First, the sea receded much farther than anyone had ever seen, exposing large areas of reefs that normally remained submerged even at low tide. People walked out onto the exposed reefs in wonder. But then, with a rumble that grew to a roar, a wall of frothing water began to build in the distance and approach land (figure above b). Puzzled bathers first watched, then ran inland in panic when the threat became clear. As the tsunami approached shore, friction with the sea floor had slowed it to less than 30 km an hour, but it still moved faster than people could run. In places, the wave front reached heights of 15 to 30 m (45 to 100 feet) as it slammed into Banda Aceh (figure above c). The impact of the water ripped boats from their moorings, snapped trees, battered buildings into rubble, and tossed cars and trucks like toys. And the water just kept coming, eventually flooding low-lying land up to 7 km inland (figure above d). It drenched forests and fields with salt water (deadly to plants) and buried fields and streets with up to a meter of sand and mud. When the water level finally returned to normal, a jumble of flotsam, as well as the bodies of unfortunate victims, were dragged out to sea and drifted away. Geologists refer to the tsunami that struck Banda Aceh as a near-field (or local) tsunami, because of its proximity to the earthquake. But the horror of Banda Aceh was merely a preamble to the devastation that would soon visit other stretches of Indian Ocean coast. Far-field (or distant) tsunamis crossed the ocean and struck Sri Lanka 2.5 hours after the earthquake, the coast of India half an hour after that, and the coast of Africa, on the west side of the Indian Ocean, 5.5 hours after the earthquake. In the end, more than 230,000 people died that day. The tsunami that struck Japan soon after the 2011 Tohoku earthquake was vividly captured in high-definition video that was seen throughout the world, generating a new level of international awareness. Though much of the coast was fringed by seawalls, they proved to be a minor impediment to the advance of the wave, which, in places, was 10 to 30 m high when it reached shore. Racing inland the wave erased whole towns, submerging airports and fields. As the wave picked up dirt and debris, it became a viscous slurry, moving with such force that nothing could withstand its impact. 

The devastation of coastal towns was so complete that they looked as though they had been struck by nuclear bombs  (figure above a). But the catastrophe was not over. The wave had also hit a nuclear power plant. Though the plant had withstood ground shaking and had automatically shut down, its radioactive core still needed to be cooled by water in order to remain safe. The tsunami not only destroyed power lines, cutting the plant off from the electrical grid, but it also eliminated backup diesel generators and cut water lines. Thus, cooling pumps stopped functioning. Eventually, water surrounding the heat producing radioactive core of the reactors, as well as the water cooling spent fuel, boiled away. Some of the water separated into hydrogen and oxygen gas, which exploded, and ultimately, the integrity of the nuclear plant was breached so that radioactivity entered the environment (figure above b). Because tsunamis are so dangerous, predicting their arrival can save thousands of lives. A tsunami warning centre in Hawaii keeps track of earthquakes around the Pacific and uses data relayed from tide gauges and sea-floor pressure gauges to determine whether a particular earthquake has generated a tsunami. If observers detect a tsunami, they flash warnings to authorities around the Pacific. 


Once the ground shaking and fires have stopped, disease may still threaten lives in an earthquake damaged region. Earthquakes cut water and sewer lines, destroying clean-water supplies and exposing the public to bacteria, and they cut transportation lines, preventing food and medicine from reaching the area. The severity of such problems depends on the ability of emergency services to cope. The lack of sufficient clean water after the 2010 Haiti earthquake led to a cholera epidemic later that year.

Can we predict Earthquakes?

Can seismologists predict earthquakes? 

The answer depends on the time frame of the prediction. With our present understanding of the distribution of seismic zones and the frequency at which earthquakes occur, we can make long-term predictions (on the time scale of decades to centuries). For example, with some certainty, we can say that a major earthquake will rattle Istanbul during the next 100 years, and that a major earthquake probably won’t strike central Canada during the next 10 years. But despite extensive research, seismologists cannot make accurate short-term predictions (on the time scale of hours to weeks or even years). Thus we cannot say, for example, that an earthquake will happen in Montreal at 2:43 P.M. on January 17. In this section, we look at the scientific basis of both long- and short-term predictions and consider the consequences of a prediction. Seismologists refer to studies leading to predictions as seismic-risk, or  seismic-hazard assessment. 

Long-Term Predictions 

A long-term prediction estimates the probability, or likelihood that an earthquake will happen during a specified time range. For example, a seismologist may say, “The probability of a major earthquake occurring in the next 20 years in this state is 20%.” This sentence implies that there’s a 1-in-5 chance that the earthquake will happen before 20 years have passed. Urban planners and civil engineers can use long-term predictions to help create building codes for a region codes requiring stronger, more expensive buildings make sense for regions with greater seismic risk. They may also use predictions to determine whether it is reasonably safe to build vulnerable structures such as nuclear power plants, hospitals, or dams in a given region. Seismologists base long-term earthquake predictions on two pieces of information: the identification of seismic zones and the recurrence interval (the average time between successive events). 

To identify a seismic zone, seismologists produce a map showing the epicentres of earthquakes that have happened  during a set period of time (say, 30 years). Clusters or belts of epicentres define the seismic zone. The basic premise of long-term earthquake prediction can be stated as follows: a region in which there have been many earthquakes in the past will be more likely to experience earthquakes in the future. Seismic zones, therefore, are regions of greater seismic risk. This doesn't mean that a disastrous earthquake can’t happen far from a seismic zone they can and do but the probability that an event will happen in a given time window is less. To determine the recurrence interval for large earthquakes within a given seismic zone, seismologists must determine when large earthquakes happened there in the past. Since the historical record does not provide information far enough back in time, they study geologic evidence for great earthquakes. For example, recognition of a fresh, unweathered fault scarp or trace may indicate that faulting affected an area relatively recently. A trench cut into sedimentary strata near a fault may reveal layers of sand volcanoes and disrupted bedding in the stratigraphic record. Each layer, whose age can be determined by using radiocarbon dating of plant fragments, records the time of an earthquake (figure above). By calculating the number of years between successive events and taking the average, seismologists obtain the recurrence interval. Note that a recurrence interval does not specify the exact number of years between events, only the average number. Since stress builds up over time on a fault, the probability that an earthquake will happen in any given year probably increases as time passes. 

Information on a recurrence interval allows seismologists to refine regional maps illustrating seismic risk (figure above a and b).

Short-Term Predictions 

Short-term predictions, specifying that an earthquake will happen on a given date or within a time window of days to years, are not and may never be reliable. Seismologists have considered, and discounted as unreliable, many supposed bases for short-term prediction. For example, a swarm of fore-shocks may indicate that rock is beginning to crack in advance of a main-shock, but such swarms can be identified only in hindsight. Precise surveys show that the surface of the ground may warp slightly prior to an earthquake, but no one can determine how much warping will take place before an earthquake will happen. Prediction studies focused on measuring changes in water levels in wells, radon gas in spring water, electrical signals emitted by minerals, or agitation of animals have met with similar skepticism. The concept of a short-term prediction should not be confused with the concept of an earthquake early warning system. An early warning system works as follows. When an earthquake happens, the seismic waves it produces start travelling through the Earth. Seismic stations closer to the epicentre may detect an earthquake before the seismic waves have had time to reach populated areas farther from the epicentre. The instant that seismic stations detect the earthquake, a computer approximates the epicentre location, then sends a signal to a control centre, which automatically sends out emergency signals to areas that might be affected. The signals shut down gas pipelines, trains, nuclear reactors, power lines, and other vulnerable infrastructure. The signal also sets off sirens and alerts broadcasters to send out warnings on radio, TV, and cell-phone networks to warn people that an earthquake is about to begin. Unless the focus is directly under the city, the warning may precede the arrival of the first earthquake waves by several seconds, not a lot of time, but hopefully enough to prevent some infrastructure damage and perhaps enough for people to seek a safer location.

Thursday, 19 November 2015

Where and Why Do Earthquakes Occur?

Earthquakes occurring place

How earthquakes happen? Where do most earthquakes occur? Why do earthquakes happen? How do earthquakes happen? Where are earthquakes most likely to occur? Why do earthquakes happen? Where in the world do most earthquakes occur?
Earthquakes do not take place everywhere on the globe. By plotting the distribution of earthquake epicentres on a map, seismologists find where do most earthquakes occur, but not all, earthquakes occur in fairly narrow seismic belts, or seismic zones. Most seismic belts correspond to plate boundaries, and where most earthquakes occur within these belts are called plate-boundary earthquakes. Earthquakes that occur away from plate boundaries are called intra-plate earthquakes (figure above). Eighty percent of the earthquake energy released on Earth comes from the plate-boundary earthquakes in the belts surrounding the Pacific Ocean. 
Earthquakes do not occur at random depths in the Earth. Seismologists distinguish three classes of earthquakes based on focus depth: shallow-focus earthquakes occur in the top 60 km of the Earth, intermediate-focus earthquakes take place between 60 and 300 km, and deep-focus earthquakes occur down to a depth of about 660 km. Earthquakes cannot happen still deeper in the Earth, because rock at great depth cannot rupture or change in a manner that produces shock waves. In this section, we look at the characteristics of earthquakes in various geologic settings and learn why earthquakes take place where they do. 

Plate-Boundary Earthquakes 

As we've noted, the most earthquakes occur at faults along plate boundaries, for the relative motion between plates causes slip on faults. We find different kinds of faulting at different types of plate boundaries.

Divergent plate-boundary seismicity

At divergent plate boundaries (mid-ocean ridges), two oceanic plates form and move apart. Divergent boundaries are broken into spreading segments linked by transform faults. Therefore, two kinds of faults develop at divergent boundaries. Along spreading segments, stretching generates normal faults, whereas along transform faults strike-slip displacement occurs (figure above). 

Transform plate-boundary seismicity

At transform plate boundaries, where one plate slides past another without the production or consumption of oceanic lithosphere, most faulting results in strike-slip motion. The majority of transform faults in the world link segments of oceanic ridges, but a few, such as the San Andreas fault of California, the Alpine fault of New Zealand, and the Anatolian faults in Turkey, cut through continental lithosphere or volcanic arcs. All transform-fault earthquakes have a shallow focus, so the larger ones on land can cause disaster. The 2010 earthquake in Haiti is a tragic example of such an earthquake (Box 8.1). As another example of a transform-fault earthquake, consider the slip of the San Andreas fault near San Francisco in 1906 (a in figure above). In the wake of the gold rush, San Francisco was a booming city with broad streets and numerous large buildings. But it was built on the transform boundary along which the Pacific Plate moves north at an average rate of 6 cm per year, relative to North America. Because of the stick-slip behaviour of the fault, this movement doesn't occur smoothly but happens in jerks, each of which causes an earthquake. At 5:12 A.M. on April 18, the fault near San Francisco slipped by as much as 7 m, and earthquake waves slammed into the city. Witnesses watched in horror as the street undulated like ocean waves. Buildings swayed and banged together, laundry lines stretched and snapped, church bells rang, and then towers, facades, and houses toppled. Judging from the damage, seismologists estimate that the largest shock would have registered a seismic moment magnitude of 7.9. Most building collapse took place down-town. Fire followed soon after, consuming huge areas of the city, for most buildings were made of wood (b in figure above). In the end, about five hundred people died, and a quarter of a million were left homeless. The San Francisco earthquake has not been the only one to strike along the San Andreas and nearby related faults. Over a dozen major earthquakes have happened on these faults during the past two centuries, including the 1857 magnitude 7.7 earthquake just east of Los Angeles, and the 1989 MW 7.1 Loma Prieta earthquake, which occurred 100 km south of San Francisco but nevertheless shut down a World Series game and caused the collapse of a double-decker freeway (c in figure above).

Convergent plate-boundary seismicity

Convergent plate boundaries are complicated regions at which several different kinds of earthquakes take place. Shallow-focus earthquakes occur in both the sub-ducting plate and the overriding  plate. Specifically, as the down-going plate begins to  subduct, it scrapes along the base of the overriding plate. Large thrust faults develop along the contact between the down-going and overriding plates, and shear on these faults can produce disastrous, shallow earthquakes. In some cases, the push applied by the down-going plate compresses the overriding plate and triggers shallow earthquakes in the overriding  plate. In contrast to other types of plate boundaries, convergent plate boundaries also host intermediate-focus and deep focus earthquakes. These occur in the down-going slab as it sinks into the mantle, defining the sloping band of seismicity called a Wadati-Benioff zone, after the seismologists who first recognised it (a in figure above). These earthquakes occur partly in response to stresses caused by shear between the down-going lithosphere plate and surrounding asthenosphere, and partly by the pull of the sinking deeper part of the plate on the shallower part. Why can intermediate- and deep-focus earthquakes of a Wadati-Benioff zone take place? Shouldn't the rock of a subducted plate at these depths be too warm and soft to break seismically? Seismologists have determined that rock is such a good insulator that the interior of a plate actually remains cool enough to break seismically, even down to a depth of about 300 km. And they found that at the extreme pressures developed in deeply subducted lithosphere, certain minerals can suddenly collapse to form new, denser minerals, a process that could generate an earthquake. Earthquakes in southern Alaska, eastern Japan, the western coast of South America, the coast of Oregon and Washington, and along island arcs in the western Pacific serve as examples of convergent-boundary earthquakes. Some of these earthquakes are large, and occur near populated areas, so they can be devastating. 
Notable examples include the 1960 MW 9.5 earthquake in Chile, the largest earthquake on record; the 1964 MW 9.2 Good Friday earthquake near Anchorage, Alaska; the 1995 MW 6.9 earthquake in Kobe, Japan, which devastated the city (b & c in figure above); the 2004 MW 9.3 Sumatra earthquake, which triggered the giant Indian Ocean waves that killed 230,000 people; the 2010 MW 8.8 Chilean earthquake; and the 2011 MW 9.0 Tohoku earthquake, which also triggered a tsunami.

Earthquakes due to Rifting and Collision

Continental rifts

The stretching of crust at continental rifts generates normal faults (figure above). Active rifts today include the East African Rift, the Basin and Range Province (mostly in Nevada, Utah, and Arizona), and the Rio Grande Rift (in New Mexico). In all these places, shallow earthquakes occur, similar in nature to the earthquakes at mid-ocean ridges. But in contrast to mid-ocean ridges, these seismic zones can be located under populated areas, and thus cause major damage.

Collision zones

Two continents collide when the oceanic lithosphere that once separated them has been completely subducted. Such collisions produce great mountain ranges such as the Alpine-Himalayan chain and caused the catastrophic 2005 earthquake in Pakistan. Though a variety of earthquakes happen in collision zones, the most common result from movement on thrust faults (see figure above). The magnitude 9.0 earthquake that occurred in Lisbon, Portugal, on All Saint’s Day, November 1, 1755, serves as an example of a collision-related thrust event. The event happened when stresses arising from the northward push of Africa against Europe caused a thrust fault beneath the Atlantic Ocean floor west of Lisbon to slip. The resulting ground shaking toppled 85% of the city’s buildings. Fires set by overturned stoves then consumed much of the wreckage. Uplift of the sea floor by the thrust movement also produced a tsunami that inundated the coast and washed away Lisbon’s harbour.  
Not only did the catastrophe destroy irreplaceable structures (including the library that housed all the records of Portuguese exploration) and countless Renaissance artworks, but it led the intelligentsia of that time to question long-held beliefs about philosophy. Influential works by Voltaire (1694–1778) address the philosophical implications of the event.

Intraplate Earthquakes

Some earthquakes occur in the interiors of plates and are not associated with plate boundaries, active rifts, or collision zones (Divergent plate-boundary seismicity figure). These intraplate earthquakes account for only about 5% of the earthquake energy released in a year. Almost all have a shallow focus. What causes intraplate earthquakes? Most seismologists favour the idea that force applied to the boundary of a plate can cause the interior of the plate to break suddenly at weak, pre-existing fault zones, some of which may have formed initially during the Precambrian. In Europe, a number of intraplate earthquakes have been recorded. For example, an earthquake with a magnitude of 4.8 hit central England, near Birmingham, in 2002. In North America, intraplate earthquakes occur in the vicinities of New Madrid, Missouri; Charleston, South Carolina; eastern Tennessee; and Montreal. 

A magnitude 7.3 earthquake occurred near Charleston in 1886, ringing church bells up and down the coast and vibrating buildings as far away as Chicago. In Charleston itself, over 90% of the buildings were damaged, and sixty people died. In 2011, a magnitude 5.9 earthquake struck central Virginia, abruptly reminding residents of the eastern United States that the region is not immune to seismicity. The tremor was felt from the Carolinas to New England. The largest intraplate earthquakes to affect the United States happened in the early 19th century, near New Madrid, which lies near the Mississippi River in southernmost Missouri. During the winter of 1811–12, three magnitude  7 to 7.4 earthquakes struck the region. Displacement of the ground surface temporarily reversed the flow of the Mississippi River and toppled cabins (a in figure above). The earthquakes resulted from slip on faults that underlie the Mississippi Valley (b in figure above). Both St. Louis, Missouri, and Memphis, Tennessee, lie close to the epicenter, so major earthquakes in the area now could be disastrous.
Credits: Stephen Marshak (Essentials of Geology)

Friday, 13 November 2015

Consequences and Causes of Metamorphism

What Is a Metamorphic Rock? 

If someone were to put a rock on a table in front of you, how would you know that it is metamorphic? First, metamorphic rocks can possess metamorphic minerals, new minerals that grow in place within the solid rock only under metamorphic temperatures and pressures. In fact, metamorphism can produce a group of minerals that together make up what geologists call a “metamorphic mineral assemblage.” And second, metamorphic rocks can have metamorphic texture defined by distinctive arrangements of mineral grains not found in other rock types. Commonly, the texture results in metamorphic foliation, due to the parallel alignment of platy minerals (such as mica) and/ or the presence of alternating light-coloured and dark-coloured layers. When metamorphic minerals and/or textures develop, a metamorphic rock becomes as different from its protolith as a butterfly is from a caterpillar. For example, metamorphism of red shale can yield a metamorphic rock consisting of aligned mica flakes and brilliant garnet crystals (a in figure above), and metamorphism of a limestone composed of cemented-together fossil fragments can yield a metamorphic rock consisting of large interlocking crystals of calcite (b in figure above). The process of forming metamorphic minerals and textures takes place very slowly it may take thousands to millions of years and it involves several processes, which sometimes occur alone and sometimes together. The most common processes are: 
  • Recrystallization, which changes the shape and size of grains without changing the identity of the mineral making up the grains (a in figure above). 
  • Phase change, which transforms one mineral into another mineral with the same composition but a different crystal structure. On an atomic scale, phase change involves the rearrangement of atoms. 
  • Metamorphic reaction, or neocrystallization (from the Greek neos, for new), which results in growth of new mineral crystals that differ from those of the protolith (b in figure above). During neocrystallization, chemical reactions digest minerals of the protolith to produce new minerals of the metamorphic rock. 
  • Pressure solution, which happens when a wet rock is squeezed more strongly in one direction than in others. Mineral grains dissolve where their surfaces are pressed against other grains, producing ions that migrate through the water to precipitate elsewhere (c in figure above). 
  • Plastic deformation, which happens when a rock is squeezed or sheared at elevated temperatures and pressures. Under these conditions, grains behave like soft plastic and change shape without breaking (d in figure above). 

Caterpillars undergo metamorphosis because of hormonal changes in their bodies. Rocks undergo metamorphism when they are subjected to heat, pressure, compression and shear, and/or very hot water. Let’s now consider the details of how these agents of metamorphism operate.

Metamorphism Due to Heating 

When you heat cake batter, the batter transforms into a new material cake. Similarly, when you heat a rock, its ingredients transform into a new material metamorphic rock. Why? Think about what happens to atoms in a mineral grain as the grain warms. Heat causes the atoms to vibrate rapidly, stretching and bending chemical bonds that lock atoms to their neighbours. If bonds stretch too far and break, atoms detach from their original neighbours, move slightly, and form new bonds with other atoms. Repetition of this process leads to rearrangement of atoms within grains, or to migration of atoms into and out of grains, a process called solid-state diffusion. As a consequence, recrystallization and/or neo-crystallization take place, enabling a metamorphic mineral  assemblage to grow in solid rock. Metamorphism takes place at temperatures between those at which diagenesis occurs and those that cause melting. Roughly speaking, this means that most metamorphic rocks you find in outcrops on continents formed at temperatures of between 250C and 850C.

Metamorphism Due to Pressure 

As you swim underwater in a swimming pool, water squeezes against you equally from all sides in other words, your body feels pressure. Pressure can cause a material to collapse inward. For example, if you pull an air-filled balloon down to a depth of 10 m in a lake, the balloon becomes significantly smaller. Pressure can have the same effect on minerals. Near the Earth’s surface, minerals with relatively open crystal structures can be stable. However, if you subject these minerals to extreme pressure, the atoms pack more closely together and denser minerals tend to form. Such transformations involve phase changes and/or neo-crystallization.

Changing Both Pressure and Temperature 

So far, we've considered changes in pressure and temperature as separate phenomena. But in the Earth, pressure and temperature change together with increasing depth. For example, at a depth of 8 km, temperature in the crust reaches about 200C and pressure reaches about 2.3 kbar. If a rock slowly becomes buried to a depth of 20 km, as can happen during mountain building, temperature in the rock increases to more than 500C, and pressure to 5.5 kbar. Experiments and calculations show that the “stability” of certain minerals (the ability of a mineral to form and survive) depends on both pressure and temperature. When pressure and temperature increase, the original mineral assemblage in a rock becomes unstable, and a new assemblage forms out of minerals that are stable. Thus, a metamorphic rock formed at 8 km does not contain the same minerals as one formed at 20 km.

Compression, Shear, and Development  of Preferred Orientation 

Imagine that you have just built a house of cards and, being in a destructive mood, you step on it. The structure collapses because the downward push you apply with your foot exceeds the push provided by air in other directions. We can say that we have subjected the cards to compression (a in figure above). Compression flattens a material (b in figure above). Shear, in contrast, moves one part of a material sideways, relative to another. If, for example, you place a deck of cards on a table, then set your hand on top of the deck and move your hand parallel to the table, you shear the deck (c in figure above). When rocks are subjected to compression and shear at elevated temperatures and pressures, they can change shape without breaking. As it changes shape, the internal texture of a rock also changes. For example, platy (pancake-shaped) grains become parallel to one another, and elongate (cigar shaped) grains align in the same direction. Both platy and elongate grains are inequant grains, meaning that the dimension of a grain is not the same in all directions; in contrast, equant grains have roughly the same dimensions in all directions (d in figure above). The alignment of inequant minerals in a rock results in a preferred orientation (e in figure above).

The Role of Hydrothermal Fluids 

Metamorphic reactions commonly take place in the presence of hydrothermal fluids (very hot-water solutions). Where does the water in hydrothermal fluids come from? Some of it was originally bonded to minerals in the protolith, for metamorphic reactions can release such water into its surroundings. Some of it may seep up into the protolith from a nearby igneous intrusion, or down from overlying groundwater reservoirs. Notably, under extremely high pressures and temperatures, the water of hydrothermal fluids is in neither gas nor liquid state, but rather is in a “supercritical” state, meaning that it has characteristics of both gas and liquid. Such hydrothermal fluids chemically react with rock; they accelerate metamorphic reactions, because atoms involved in the reactions can migrate faster through a fluid than they can through a solid, and hydrothermal fluids provide water that can be absorbed by minerals during metamorphic reactions. Finally, fluids passing through a rock may pick up some dissolved ions and drop off others, as a bus picks up and drops off passengers, and thus can change the overall chemical composition of a rock during metamorphism. The process of changing a rock’s chemical composition by reactions with hydrothermal fluids is called metasomatism.

Tuesday, 10 November 2015

Recognizing Depositional Environments

How Do We Recognize Depositional Environments? 

Geologists refer to the conditions in which sediment was deposited as the depositional environment. Examples include beach, glacial, and river environments. To identify depositional environments, geologists, like crime scene investigators, look for clues. Detectives may seek fingerprints and bloodstains to identify a culprit. Geologists examine grain size, composition, sorting, bed-surface marks, cross bedding, and fossils to identify a depositional environment. Geological clues can tell us if the sediment was deposited by ice, strong currents, waves, or quiet water, and in some cases can provide insight into the climate at the time of deposition. With experience, geologists can examine a succession of beds and determine if it accumulated on a river floodplain, along a beach, in shallow water just offshore, or on the deep ocean floor.
Let’s now explore some examples of different depositional environments and the sediments deposited in them, by imagining that we are taking a journey from the mountains to the sea, examining sediments as we go. We will see that geologists distinguish among three basic categories of depositional environments: terrestrial, coastal, and marine.

Terrestrial (Nonmarine) Sedimentary Environments 

We begin our exploration with terrestrial depositional environments, those that develop inland, far enough away from the shoreline that they are not affected by ocean tides and waves. The sediments settle on dry land, or under and adjacent to freshwater.  
In some settings, oxygen in surface water or groundwater reacts with iron to produce rust-like iron-oxide minerals in terrestrial sediments, which give the sediment an overall reddish hue. Strata with this hue are informally called redbeds.

Glacial environments 

High in the mountains, where it’s so cold that more snow collects in the winter than melts away,  glaciers rivers or sheets of ice develop and slowly flow. Because ice is a solid, it can move sediment of any size. So as a glacier moves down a valley in the mountains, it carries along all the sediment that falls on its surface from adjacent cliffs or gets plucked from the ground at its base or sides. At the end of the glacier, where the ice finally melts away, the sediment that had been in or on the ice accumulates as “glacial till” (a in figure above). Till is unsorted and unstratified it contains clasts ranging from clay size to boulder size all mixed together.

Mountain stream environments

As we walk down beyond the end of the glacier, we enter a realm where turbulent streams rush downslope in steep-sided valleys. This fast-moving water has the power to carry large clasts; in fact, during floods, boulders and cobbles can tumble down the stream bed. Between floods, when water flow slows, the largest clasts settle out to form gravel and boulder beds, while the stream carries finer sediments like sand and mud farther downstream (b in figure above). Sedimentary deposits of a mountain stream would, therefore, include breccia and  conglomerate.

Alluvial-fan environments

Our journey now takes us to the mountain front, where the fast-moving stream empties onto a plain. In arid regions, where there is not enough water for the stream to flow continuously, the stream deposits its load of sediment near the mountain front, producing a wedge-shaped apron of gravel and sand called an alluvial fan  (c in figure above). Deposition takes place here because when the stream pours from a canyon mouth and spreads out over a broader region, friction with the ground causes the water to slow down, and slow-moving water does not have the power to move coarse sediment. The sand here still contains feldspar grains, for these have not yet weathered into clay. Alluvial-fan sediments become arkose and conglomerate.

Sand-dune environments

If the climate is very dry, few plants can grow and the ground surface lies exposed. Strong winds can move dust and sand. The dust gets carried away, and the resulting well-sorted sand can accumulate in dunes. Thus, thick layers of well-sorted sandstone, in which we can find large cross beds, are relicts of desert sand-dune environments (d in figure above).

River (fluvial) environments

In climates where streams flow, we find several distinctive depositional environments. Rivers transport gravel, sand, silt, and mud. The coarser sediments tumble along the bed in the river’s channel and collect in cross-bedded, rippled layers while the finer sediments drift along, suspended in the water. This fine sediment settles out along the banks of the river, or on the floodplain, the flat land on either side of the river that is covered with water only during floods. On the floodplain, mud layers dry out between floods, leading to the formation of mud cracks. River sediments lithify to form sandstone, siltstone, and shale. Typically, the coarser sediments of channels are surrounded by layers of fine-grained floodplain deposits, so in cross section, the channel has a lens-like shape (e in figure above). Geologists commonly refer to river deposits as fluvial sediments, from the Latin word fluvius, for river.

Lake environments

In temperate climates, where water remains at the surface throughout the year, lakes form. In lakes, the relatively quiet water can’t move coarse sediment; any coarse sediment brought into the lake by a stream settles out at the stream’s outlet. Only fine clay makes it out into the centre of the lake, where it settles to form mud on the lake bed. Thus, lake sediments typically consist of finely  laminated shale (f in figure above). 

At the mouths of streams that empty into lakes, small deltas may form. A delta is a wedge of sediment that accumulates where moving water enters standing water. Deltas were so named because the map shape of some deltas resembles the Greek letter delta ($), as we discuss further in Chapter 14. In 1885, an American geologist named G. K. Gilbert showed that such deltas contain three components (figure above): topset beds composed of gravel, foreset beds of gravel and sand, and silty bottomset beds.

Coastal and Marine Environments 

Along the seashore, a variety of distinct coastal environments occur; the character of each reflects the nature of the sediment supply and the climate. Marine environments start at the high-tide line and extend offshore, to include the deep ocean floor. The type of sediment deposited at a location depends on the climate, water depth, and whether or not clastic grains are available.

Marine delta deposits

After following the river downstream for a long distance, we reach its mouth, where it empties into the sea. Here, the river builds a delta of sediment out into the sea. River water stops flowing when it enters the sea, so sediment settles out. Large deltas are much more complex than the lake examples that Gilbert studied, for they include many different sedimentary environments including swamps, channels, floodplains, and submarine slopes. Sea-level changes may cause the positions of the different environments to move with time. Thus, deposits of an ocean-margin delta produce a great variety of sedimentary rock types (a in figure above).

Coastal beach sands

Now we leave the delta and wander along the coast. Oceanic currents transport sand along the coastline. The sand washes back and forth in the surf, so it becomes well sorted (waves winnow out silt and clay) and well rounded, and because of the back-and-forth movement of ocean water over the sand, the sand surface may become rippled (b in figure above). Thus, if you find well-sorted, medium grained sandstone, perhaps with ripple marks, you may be looking at the remnants of a beach environment.

Shallow-marine clastic deposits

From the beach, we proceed offshore. In deeper water, where wave energy does not stir the sea floor, finer sediment can accumulate. Because the water here may be only meters to a few tens of meters deep, geologists refer to this depositional setting as a shallow-marine environment. Clastic sedimentary layers that accumulate in this environment tend to be fine-grained, well-sorted, well rounded silt, and they are inhabited by a great variety of organisms such as mollusks and worms. Thus, if you see beds of siltstone and mudstone containing marine fossils, you may be looking at shallow-marine clastic deposits.

Shallow-water carbonate environments

In shallow marine settings relatively free of clastic sediment, warm, clear, nutrient-rich water hosts an abundance of organisms. Their shells, which consist of carbonate minerals, make up most of the sediment that accumulates (a and b in figure above). The nature of carbonate sediment depends on the water depth. Beaches collect sand composed of shell fragments; lagoons (protected bodies of quiet water) are sites where carbonate mud accumulates; and reefs consist of coral and coral debris. Farther offshore of a reef, we can find a sloping apron of reef fragments. Shallow-water carbonate environments transform into various kinds of limestone.

Deep-marine deposits

We conclude our journey by sailing far offshore. Along the transition between coastal regions and the deep ocean, turbidity currents deposit graded beds. In the deep-ocean realm, only fine clay and plankton provide a source for sediment. The clay eventually settles out onto the deep-sea floor, forming deposits of finely laminated mudstones, and plankton shells settle to form chalk (from calcite shells; a and b in figure above) or chert (from siliceous shells). Thus, deposits of mudstone, chalk, or bedded chert indicate a deepmarine origin.