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12 Jun 2020

June 12, 2020 Comment

What Are Earth Layers Made Of?

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).

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.

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)

11 Jun 2020

June 11, 2020 Comment

The Discovery of Sea-Floor Spreading

Military needs during World War II gave a boost to sea-ﬂoor exploration, for as submarine ﬂeets grew, navies required detailed information about bathymetry, or depth variations. The invention of echo sounding (sonar) permitted such information to be gathered quickly. Echo sounding works on the same principle that a bat uses to navigate and ﬁnd insects. A sound pulse emitted from a ship travels down through the water, bounces off the sea ﬂoor, and returns up as an echo through the water to a receiver on the ship. Since sound waves travel at a known velocity, the time between the sound emission and the echo detection indicates the distance between the ship and the sea ﬂoor. (Recall that velocity distance/time, so distance velocity s time.) As the ship travels, observers can obtain a continuous record of the depth of the sea ﬂoor. The resulting cross section showing depth plotted against location is called a bathymetric proﬁle (figure above a, b). By cruising back and forth across the ocean many times, investigators obtained a series of bathymetric proﬁles and from these constructed maps of the sea ﬂoor. (Geologists can now produce such maps much more rapidly using satellite data.) Bathymetric maps reveal several important features.

Other bathymetric features of the ocean ground.

Mid-ocean ridges: The floor beneath all major oceans includes abyssal plains, which are broad, relatively flat regions of the ocean that lie at a depth of about 4 to 5 km below sea level; and mid-ocean ridges, submarine mountain ranges whose peaks lie only about 2 to 2.5 km below sea level (figure above a). Geologists call the crest of the mid-ocean ridge the ridge axis. All mid-ocean ridges are roughly symmetrical bathymetry on one side of the axis is nearly a mirror image of bathymetry on the other side.
Deep-ocean trenches: Along much of the perimeter of the Pacific Ocean, and in a few other localities as well, the ocean floor reaches depths of 8 to 12 km deep enough to swallow Mt. Everest. These deep areas occur in elongate troughs that are now referred to as trenches (figure above b). Trenches border volcanic arcs, curving chains of active volcanoes.
Seamount chains: Numerous volcanic islands poke up from the ocean floor: for example, the Hawaiian Islands lie in the middle of the Pacific. In addition to islands that rise above sea level, sonar has detected many seamounts (isolated submarine mountains), which were once volcanoes but no longer erupt. Volcanic islands and seamounts typically occur in chains, but in contrast to the volcanic arcs that border deep ocean trenches, only one island at the end of a seamount and island chain remains capable of erupting volcanically today.
Fracture zones: Surveys reveal that the ocean floor is diced up by narrow bands of vertical cracks and broken-up rock. These fracture zones lie roughly at right angles to mid-ocean ridges. The ridge axis typically steps sideways when it intersects with a fracture zone.

New Observations on the Nature of Oceanic Crust

By the mid-20th century, geologists had observed many crucial traits of the ocean-?Oor crust. These discoveries led them to comprehend that oceanic crust differs from continental crust, and that bathymetric functions of the ocean ?Oor provide clues to the starting place of the crust. Speci?Cally:

A layer of sediment composed of clay and the tiny shells of dead plankton covers much of the ocean floor. This layer becomes progressively thicker away from the mid-ocean ridge axis. But even at its thickest, the sediment layer is too thin to have been accumulating for the entirety of Earth history.
By dredging up samples, geologists learned that oceanic crust is fundamentally different in composition from continental crust. Beneath its sediment cover, oceanic crust bedrock consists primarily of basalt it does not display the great variety of rock types found on continents.
Heat flow, the rate at which heat rises from the Earth’s interior up through the crust, is not the same everywhere in the oceans. Rather, more heat rises beneath mid-ocean ridges than elsewhere. This observation led researchers to speculate that hot magma might be rising into the crust just below the mid-ocean ridge axis.
When maps showing the distribution of earthquakes in oceanic regions became available in the years after World War II, it became clear that earthquakes do not occur randomly, but rather define distinct belts (figure below). Some belts follow trenches, some follow mid-ocean ridge axes, and others lie along portions of fracture zones. Since earthquakes define locations where rocks break and move, geologists realized that these bathymetric features are places where motion is taking place.

A 1953 map displaying the distribution of earthquake locations within the ocean basins. Note that earthquakes arise in belts.

Harry Hess and His ?Essay in Geopoetry?

Harry Hess's primary concept of sea-ground spreading. Hess implied, incorrectly, that only the crust moved. We will see that this cartoon is an oversimplified.

In the late 1950s, Harry Hess, after studying the observations described above, realized that because the sediment layer on the ocean ﬂoor was thin overall, the ocean ﬂoor might be much younger than the continents. Also, because the sediment thickened progressively away from mid-ocean ridges, the ridges themselves likely were younger than the deeper parts of the ocean floor. If this was so, then somehow new ocean floor must be forming at the ridges, and thus an ocean basin could be getting wider with time. But how? The association of earthquakes with mid-ocean ridges suggested to him that the sea ﬂoor was cracking and splitting apart at the ridge. The discovery of high heat ﬂow along mid-ocean ridge axes provided the ﬁnal piece of the puzzle, for it suggested the presence of very hot molten rock beneath the ridges. In 1960, Hess suggested that indeed molten rock (basaltic magma) rose upward beneath mid-ocean ridges and that this material solidiﬁed to form oceanic crust basalt (figure above). The new sea ﬂoor then moved away from the ridge, a process we now call sea-ﬂoor spreading. Hess realized that old ocean ﬂoor must be consumed somewhere, or the Earth would have to be expanding, so he suggested that deep ocean trenches might be places where the sea ﬂoor sank back into the mantle. Hess suggested that earthquakes at trenches were evidence of this movement, but he didn’t understand how the movement took place. Other geologists, such as Robert Dietz, were coming to similar conclusions at about the same time.

Hess and his contemporaries realized that the ocean-?Oor-spreading speculation immediately provided the long sought clarification of ways continental ?Drift? Occurs. Continents passively move aside as the ocean ?Oor among them spreads at mid-ocean ridges, and they passively pass together as the sea ?Oor between them sinks returned into the mantle at trenches. (As we will see later, geologists now comprehend that it's miles the lithosphere that movements, now not simply the crust.) Thus, sea-?Oor spreading proved to be an important step on the direction to plate tectonics the idea seemed so exact that Hess cited his description of it as ?An essay in geopoetry.? But ?Rst, the idea needed to be examined, and other key discoveries would have to take area earlier than the entire concept of plate tectonics should come together.

June 11, 2020 Comment

Evidence for Sea-Floor Spreading

Marine Magnetic Anomalies

Recognizing anomalies

For a hypothesis to become a theory, researchers must demonstrate that the idea really works. During the 1960s, geologists found that the sea-ﬂoor spreading hypothesis successfully explains several previously bafﬂing observations. Here we discuss two: (1) the existence of orderly variations in the strength of the measured magnetic ﬁeld over the sea ﬂoor, producing a pattern of stripes called marine magnetic anomalies; and (2) the variation in sediment thickness on the ocean crust, as measured by drilling.

Geologists can measure the strength of Earth’s magnetic ﬁeld with an instrument called a magnetometer. At any given location on the surface of the Earth, the magnetic ﬁeld that you measure includes two parts: one produced by the main dipole of the Earth generated by circulation of molten iron in the outer core, and another produced by the magnetism of near-surface rock. A magnetic anomaly is the difference between the expected strength of the Earth’s main dipole ﬁeld at a certain location and the actual measured strength of the magnetic ﬁeld at that location. Places where the ﬁeld strength is stronger than expected are positive anomalies, and places where the ﬁeld strength is weaker than expected are negative anomalies.

The discovery of marine magnetic anomalies.

Geologists towed magnetometers back and forth across the ocean to map variations in magnetic ﬁeld strength (figure above a). As a ship cruised along its course, the magnetometer’s gauge might ﬁrst detect an interval of strong signal (a positive anomaly) and then an interval of weak signal (a negative anomaly). A graph of signal strength versus distance along the traverse, therefore, has a sawtooth shape (figure above b). When geologists compiled data from many cruises on a map, these marine magnetic anomalies deﬁned distinctive, alternating bands. If we color positive anomalies dark and negative anomalies light, the pattern made by the anomalies resembles the stripes on a candy cane (figure above c). The mystery of this marine magnetic anomaly pattern, however, remained unsolved until geologists recognized the existence of magnetic reversals.

Recall that Earth’s magnetic ﬁeld can be represented by an arrow, representing the dipole, that presently points from the north magnetic pole to the south magnetic pole. When researchers measured the paleomagnetism of a succession of rock layers that had accumulated over a long period of time, they found that the polarity (which end of a magnet points north and which end points south) of the paleomagnetic ﬁeld preserved in some layers was the same as that of Earth’s present magnetic ﬁeld, whereas in other layers it was the opposite (figure above a, b).

At ﬁrst, observations of reversed polarity were largely ignored, thought to be the result of lightning strikes or of local chemical reactions between rock and water. But when repeated measurements from around the world revealed a systematic pattern of alternating normal and reversed polarity in rock layers, geologists realized that reversals were a worldwide, not a local, phenomenon. They reached the unavoidable conclusion that, at various times during Earth history, the polarity of Earth’s magnetic ﬁeld has suddenly reversed! In other words, sometimes the Earth has normal polarity, as it does today, and sometimes it has reversed polarity (figure above c). A time when the Earth’s ﬁeld ﬂips from normal to reversed polarity, or vice versa, is called a magnetic reversal. When the Earth has reversed polarity, the south magnetic pole lies near the north geographic pole, and the north magnetic pole lies near the south geographic pole. Thus, if you were to use a compass during periods when the Earth’s magnetic ﬁeld was reversed, the north-seeking end of the needle would point to the south geographic pole. Note that the Earth itself doesn't turn upside down it is just the magnetic ﬁeld that reverses.

In the 1950s, about the same time researchers discovered polarity reversals, they developed a technique that permitted them to measure the age of a rock in years. Geologists applied the technique to determine the ages of rock layers in which they obtained their paleomagnetic measurements, and thus determined when the magnetic ﬁeld of the Earth reversed. With this information, they constructed a history of magnetic reversals for the past 4.5 million years; this history is now called the magnetic-reversal chronology. The time interval between successive reversals is called a chron.

A diagram representing the Earth’s magnetic-reversal chronology (figure above d) shows that reversals do not occur regularly, so the lengths of different polarity chrons are different. For example, we have had a normal-polarity chron for about the last 700,000 years. Before that, a reversed-polarity chron occurred. The youngest four polarity chrons (Brunhes, Matuyama, Gauss, and Gilbert) were named after scientists who had made important contributions to the study of magnetism. As more measurements became available, investigators realized that some short-duration reversals (less than 200,000 years long) took place within the chrons, and they called these shorter durations “polarity subchrons.” Using isotopic dating, it was possible to determine the age of chrons back to 4.5 Ma.

Interpreting marine magnetic anomalies

The progressive development of magnetic anomalies and the long-term reversals chronology.

Why do marine magnetic anomalies exist? In 1963, researchers in Britain and Canada proposed a solution to this riddle. Simply put, a positive anomaly occurs over areas of the sea ﬂoor where underlying basalt has normal polarity. In these areas, the magnetic force produced by the magnetite grains in basalt adds to the force produced by the Earth’s dipole the sum of these forces yields a stronger magnetic signal than expected due to the dipole alone (figure above a). A negative anomaly occurs over regions of the sea ﬂoor where the underlying basalt has a reversed polarity. In these regions, the magnetic force of the basalt subtracts from the force produced by the Earth’s dipole, so the measured magnetic signal is weaker than expected.

The sea-ﬂoor-spreading model easily explains not only why positive and negative magnetic anomalies exist over the sea ﬂoor, but also why they deﬁne stripes that trend parallel to the mid-ocean ridge and why the pattern of stripes on one side of the ridge is the mirror image of the pattern on the other side (figure above b). To see why, let’s examine stages in the process of sea-ﬂoor spreading (figure above c). Imagine that at Time 1 in the past, the Earth’s magnetic ﬁeld has normal polarity. As the basalt rising at the mid-ocean ridge during this time interval cools and solidiﬁes, the tiny magnetic grains in basalt align with the Earth’s ﬁeld, and thus the rock as a whole has a normal polarity. Sea ﬂoor formed during Time 1 will therefore generate a positive anomaly and appear as a dark stripe on an anomaly map. As it forms, the rock of this stripe moves away from the ridge axis, so half goes to the right and half to the left. Now imagine that later, at Time 2, Earth’s ﬁeld has reversed polarity. Sea-ﬂoor basalt formed during Time 2, therefore, has reversed polarity and will appear as a light stripe on an anomaly map. As it forms, this reversed-polarity stripe moves away from the ridge axis, and even younger crust forms along the axis. The basalt in each new stripe of crust preserves the polarity that was present at the time it formed, so as the Earth’s magnetic ﬁeld ﬂips back and forth, alternating positive and negative anomaly stripes form. A positive anomaly exists over the ridge axis today because sea ﬂoor is forming during the present chron of normal polarity.

Closer examination of a sea-ﬂoor magnetic anomaly map reveals that anomalies are not all the same width. Geologists found that the relative widths of anomaly stripes near the Mid-Atlantic Ridge are the same as the relative durations of paleomagnetic chrons (figure above d). This relationship between anomaly-stripe width and polarity-chron duration indicates that the rate of sea-ﬂoor spreading has been constant along the Mid-Atlantic Ridge for at least the last 4.5 million years. If you assume that the spreading rate was constant for tens to hundreds of millions of years, then it is possible to estimate the age of stripes right up to the edge of the ocean.

Evidence from Deep-Sea Drilling

In the late 1960s, a research drilling ship called the Glomar Challenger set out to sail around the ocean drilling holes into the sea ﬂoor. This amazing ship could lower enough drill pipe to drill in 5-km-deep water and could continue to drill until the hole reached a depth of about 1.7 km (1.1 miles) below the sea ﬂoor. Drillers brought up cores of rock and sediment that geoscientists then studied on board.

On one of its early cruises, the Glomar Challenger drilled a series of holes through sea-ﬂoor sediment to the basalt layer. These holes were spaced at progressively greater distances from the axis of the Mid-Atlantic Ridge. If the model of sea-ﬂoor spreading was correct, then not only should the sediment layer be progressively thicker away from the axis, but the age of the oldest sediment just above the basalt should be progressively older away from the axis. When the drilling and the analyses were complete, the prediction was conﬁrmed. Thus, studies of both marine magnetic anomalies and the age of the sea ﬂoor proved the sea-ﬂoor-spreading model.

Credits: Stephen Marshak ( Essentials of Geology)

June 11, 2020 Comment

What Do We Mean by Plate Tectonics?

The paleomagnetic proof of continental drift (plate tectonics) and the discovery of sea-ﬂoor spreading set off a scientiﬁc revolution in geology in the 1960s and 1970s. Geologists realised that many of their existing interpretations of global geology, based on the premise that the positions of continents and oceans remain ﬁxed in position through time, were simply wrong! Researchers dropped what they were doing and turned their attention to studying the broader implications of continental drift and sea-ﬂoor spreading. It became clear that these phenomena required that the outer shell of the Earth was divided into rigid plates that moved relative to each other. New studies clariﬁed the meaning of a plate, deﬁned the types of plate boundaries, constrained plate motions, related plate motions to earthquakes and volcanoes, showed how plate interactions can explain mountain belts and seamount chains, and outlined the history of past plate motions. From these, the modern theory of plate tectonics evolved. Below, we ﬁrst describe lithosphere plates and their boundaries, and then outline the basic principles of plate tectonics theory.

We learned earlier that geoscientists divide the outer part of the Earth into two layers. The lithosphere consists of the crust plus the top (cooler) part of the upper mantle. It behaves relatively rigidly, meaning that when a force pushes or pulls on it, it does not ﬂow but rather bends or breaks (figure above a). The lithosphere ﬂoats on a relatively soft, or “plastic,” layer called the asthenosphere, composed of warmer ( 1280°C) mantle that can ﬂow slowly when acted on by a force. As a result, the asthenosphere convects, like water in a pot, though much more slowly.

Continental lithosphere and oceanic lithosphere differ markedly in their thicknesses. On average, continental lithosphere has a thickness of 150 km, whereas old oceanic lithosphere has a thickness of about 100 km (figure above b). (For reasons discussed later in this chapter, new oceanic lithosphere at a mid-ocean ridge is much thinner.) Recall that the crustal part of continental lithosphere ranges from 25 to 70 km thick and consists largely of low-density felsic and intermediate rock. In contrast, the crustal part of oceanic lithosphere is only 7 to 10 km thick and consists largely of relatively high-density maﬁc rock (basalt and gabbro). The mantle part of both continental and oceanic lithosphere consists of very high-density ultramaﬁc rock (peridotite). Because of these differences, the continental lithosphere “ﬂoats” at a higher level than does the oceanic lithosphere.

The area of plate boundaries and the distribution of earthquakes.

The lithosphere forms the Earth’s relatively rigid shell. But unlike the shell of a hen’s egg, the lithospheric shell contains a number of major breaks, which separate it into distinct pieces. As noted earlier, we call the pieces lithosphere plates, or simply plates. The breaks between plates are known as plate boundaries (figure above a). Geoscientists distinguish twelve major plates and several microplates.

The Basic Principles of Plate Tectonics

With the historical past provided above, we will restate plate tectonics theory concisely as follows. The Earth?S lithosphere is split into plates that pass relative to each different. As a plate movements, its internal area remains broadly speaking, however now not perfectly, rigid and intact. But rock alongside plate limitations undergoes extreme deformation (cracking, sliding, bending, stretching, and squashing) because the plate grinds or scrapes towards its neighbours or pulls faraway from its neighbours. As plates circulate, so do the continents that shape a part of the plates. Because of plate tectonics, the map of Earth?S surface continuously adjustments.

Identifying Plate Boundaries

How do we recognize the location of a plate boundary? The answer becomes clear from looking at a map showing the locations of earthquakes (figure above b). Recall from Chapter 1 that earthquakes are vibrations caused by shock waves that are generated where rock breaks and suddenly slips along a fault. The epicentre marks the point on the Earth’s surface directly above the earthquake. Earthquake epicentres do not speckle the globe randomly, like buckshot on a target. Rather, the majority occur in relatively narrow, distinct belts. These earthquake belts deﬁne the position of plate boundaries because the fracturing and slipping that occurs along plate boundaries generates earthquakes. Plate interiors, regions away from the plate boundaries, remain relatively earthquake-free because they do not accommodate as much movement. While earthquakes serve as the most deﬁnitive indicator of a plate boundary, other prominent geologic features also develop along plate boundaries.

Note that some plates consist entirely of oceanic lithosphere, whereas some plates consist of both oceanic and continental lithosphere. Also, note that not all plates are the same size (figure above c). Some plate boundaries follow continental margins, the boundary between a continent and an ocean, but others do not. For this reason, we distinguish between active margins, which are plate boundaries, and passive margins, which are not plate boundaries. Earthquakes are common at active margins, but not at passive margins. Along passive margins, continental crust is thinner than in continental interiors. Thick (10 to 15 km) accumulations of sediment cover this thinned crust. The surface of this sediment layer is a broad, shallow (less than 500 m deep) region called the continental shelf, home to the major ﬁsheries of the world.

The three kinds of plate barriers range primarily based on the character of relative motion.

Geologists de?Ne 3 types of plate boundaries, based clearly on the relative motions of the plates on both side of the boundary (discern above a?C). A boundary at which plates flow other than every other is a divergent boundary. A boundary at which two plates move in the direction of each other in order that one plate sinks below the opposite is a convergent boundary. And a boundary at which two plates slide sideways beyond every different is a remodel boundary.

Credits: Stephen Marshak (Essentials of Geology)

10 Jun 2020

June 10, 2020 Comment

Convergent Plate Boundaries and Subduction

At convergent plate boundaries, two plates, at least one of which is oceanic, move toward one another. But rather than butting each other like angry rams, one oceanic plate bends and sinks down into the asthenosphere beneath the other plate. Geologists refer to the sinking process as subduction, so convergent boundaries are also known as subduction zones. Because subduction at a convergent boundary consumes old ocean lithosphere and thus ‘‘consumes’’ oceanic basins, geologists also refer to convergent boundaries as consuming boundaries, and because they are delineated by deep-ocean trenches, they are sometimes simply called trenches. The amount of oceanic plate consumption worldwide, averaged over time, equals the amount of sea-ﬂoor spreading worldwide, so the surface area of the Earth remains constant through time.

During the process of subduction, oceanic lithosphere sinks back into the deeper mantle.

Subduction occurs for a simple reason: oceanic lithosphere, once it has aged at least 10 million years, is denser than the underlying asthenosphere and thus can sink through the asthenosphere if given an opportunity. Where it lies ﬂat on the surface of the asthenosphere, oceanic lithosphere can’t sink. However, once the end of the convergent plate bends down and slips into the mantle, it continues downward like an anchor falling to the bottom of a lake (figure above a). As the lithosphere sinks, asthenosphere ﬂows out of its way, just as water ﬂows out of the way of a sinking anchor. But unlike water, the asthenosphere can ﬂow only very slowly, so oceanic lithosphere can sink only very slowly, at a rate of less than about 15 cm per year. To visualize the difference, imagine how much faster a coin can sink through water than it can through honey.

Note that the “downgoing plate,” the plate that has been subducted, must be composed of oceanic lithosphere. The overriding plate, which does not sink, can consist of either oceanic or continental lithosphere. Continental crust cannot be subducted because it is too buoyant; the low-density rocks of continental crust act like a life preserver keeping the continent aﬂoat. If continental crust moves into a convergent margin, subduction eventually stops. Because of subduction, all ocean ﬂoor on the planet is less than about 200 million years old. Because continental crust cannot subduct, some continental crust has persisted at the surface of the Earth for over 3.8 billion years.

Earthquakes and the Fate of Subducted Plates

At convergent plate boundaries, the downgoing plate grinds along the base of the overriding plate, a process that generates large earthquakes. These earthquakes occur fairly close to the Earth’s surface, so some of them cause massive destruction in coastal cities. But earthquakes also happen in downgoing plates at greater depths. In fact, geologists have detected earthquakes within downgoing plates to a depth of 660 km. The band of earthquakes in a downgoing plate is called a WadatiBenioff zone, after its two discoverers (figure above b).

At depths greater than 660 km, conditions leading to earthquakes in subducted lithosphere evidently do not occur. Recent observations, however, indicate that some downgoing plates do continue to sink below a depth of 660 km they just do so without generating earthquakes. In fact, the lower mantle may be a graveyard for old subducted plates.

Geologic Features of a Convergent Boundary

To become familiar with the various geologic features that occur along a convergent plate boundary, let’s look at an example, the boundary between the western coast of the South American Plate and the eastern edge of the Nazca Plate (a portion of the Paciﬁc Ocean ﬂoor). A deep-ocean trench, the Peru-Chile Trench, delineates this boundary (figure above b). Such trenches form where the plate bends as it starts to sink into the asthenosphere.

In the Peru-Chile Trench, as the downgoing plate slides under the overriding plate, sediment (clay and plankton) that had settled on the surface of the downgoing plate, as well as sand that fell into the trench from the shores of South America, gets scraped up and incorporated in a wedge-shaped mass known as an accretionary prism (figure above c). An accretionary prism forms in basically the same way as a pile of snow or sand in front of a plow, and like snow, the sediment tends to be squashed and contorted.

A chain of volcanoes known as a volcanic arc d evelops behind the accretionary prism. The magma that feeds these volcanoes forms just above the surface of the downgoing plate where the plate reaches a depth of about 150 km below the Earth’s surface. If the volcanic arc forms where an oceanic plate subducts beneath continental lithosphere, the resulting chain of volcanoes grows on the continent and forms a continental volcanic arc. (In some cases, the plates squeeze together across a continental arc, causing a belt of faults to form behind the arc.) If, however, the volcanic arc grows where one oceanic plate subducts beneath another oceanic plate, the resulting volcanoes form a chain of islands known as a volcanic island arc (figure above d). A back-arc basin exists either where subduction happens to begin offshore, trapping ocean lithosphere behind the arc, or where stretching of the lithosphere behind the arc leads to the formation of a small spreading ridge behind the arc (figure above e).

Credits: Stephen Marshak (Essentials of Geology)

June 10, 2020 Comment

Transform Plate Boundaries

When researchers commenced to explore the bathymetry of midocean ridges in detail, they observed that mid-ocean ridges aren't lengthy, uninterrupted traces, but rather consist of brief segments that appear to be offset laterally from each different (parent above a) by means of narrow belts of broken and abnormal sea ?Oor. These belts, or fracture zones, lie kind of at proper angles to the ridge segments, intersect the ends of the segments, and increase past the ends of the segments. Originally, researchers incorrectly assumed that the complete period of every fracture quarter changed into a fault, and that slip on a fracture region had displaced segments of the mid-ocean ridge sideways, relative to every different. In different phrases, they imagined that a mid-ocean ridge initiated as a continuous, fence-like line that most effective later become damaged up via faulting. But while information approximately the distribution of earthquakes alongside mid-ocean ridges became available, it became clean that this model couldn't be correct. Earthquakes, and consequently lively fault slip, arise best at the section of a fracture region that lies between two ridge segments. The quantities of fracture zones that expand beyond the edges of ridge segments, out into the abyssal simple, aren't seismically active.

The distribution of movement along fracture zones remained a mystery until a Canadian researcher, J. Tuzo Wilson, began to think about fracture zones in the context of the sea-ﬂoor-spreading concept. Wilson proposed that fracture zones formed at the same time as the ridge axis itself, and thus the ridge consisted of separate segments to start with. These segments were linked (not offset) by fracture zones. With this idea in mind, he drew a sketch map showing two ridge-axis segments linked by a fracture zone, and he drew arrows to indicate the direction that ocean ﬂoor was moving, relative to the ridge axis, as a result of sea-ﬂoor-spreading (figure above b). Look at the arrows in figure above b. Clearly, the movement direction on the active portion of the fracture zone must be opposite to the movement direction that researchers originally thought occurred on the structure. Further, in Wilson’s model, slip occurs only along the segment of the fracture zone between the two ridge segments (figure above c). Plates on opposite sides of the inactive part of a fracture zone move together, as one plate.

Wilson delivered the time period rework boundary, or transform fault, for the actively slipping segment of a fracture area between two ridge segments, and he pointed out that these are a third form of plate boundary. At a remodel boundary, one plate slides sideways past another, but no new plate forms and no vintage plate is fed on. Transform obstacles are, therefore, de?Ned by means of a vertical fault on which the slip route parallels the Earth?S surface. The slip breaks up the crust and bureaucracy a set of steep fractures.

So far we've discussed only transforms along mid-ocean ridges. Not all transforms link ridge segments. Some, such as the Alpine Fault of New Zealand, link trenches, while others link a trench to a ridge segment. Further, not all transform faults occur in oceanic lithosphere; a few cut across continental lithosphere. The San Andreas Fault, for example, which cuts across California, deﬁnes part of the plate boundary between the North American Plate and the Paciﬁc Plate the portion of California that lies to the west of the fault (including Los Angeles) is part of the Paciﬁc Plate, while the portion that lies to the east of the fault is part of the North American Plate (figure above d, e).

Credits: Stephen Marshak (Essentials of Geology)

June 10, 2020 Comment

How Do Plate Boundaries Form and Die?

The conﬁguration of plates and plate boundaries visible on our planet today has not existed for all of geologic history, and will not exist indeﬁnitely into the future. Because of plate motion, oceanic plates form and are later consumed, while continents merge and later split apart. How does a new divergent boundary come into existence, and how does an existing convergent boundary eventually cease to exist? Most new divergent boundaries form when a continent splits and separates into two continents. We call this process rifting. A convergent boundary ceases to exist when a piece of buoyant lithosphere, such as a continent or an island arc, moves into the subduction zone and, in effect, jams up the system. We call this process collision.

Continental Rifting

During the process of rifting, lithosphere stretches.

A continental rift is a linear belt in which continental lithosphere pulls apart (figure above a). During the process, the lithosphere stretches horizontally and thins vertically, much like a piece of taffy you pull between your ﬁngers. Nearer the surface of the continent, where the crust is cold and brittle, stretching causes rock to break and faults to develop. Blocks of rock slip down the fault surfaces, leading to the formation of a low area that gradually becomes buried by sediment. Deeper in the crust, and in the underlying lithospheric mantle, rock is warmer and softer, so stretching takes place in a plastic manner without breaking the rock. The whole region that stretches is the rift, and the process of stretching is called rifting.

As continental lithosphere thins, hot asthenosphere rises beneath the rift and starts to melt. Eruption of the molten rock produces volcanoes along the rift. If rifting continues for a long enough time, the continent breaks in two, a new midocean ridge forms, and sea-ﬂoor spreading begins. The relict of the rift evolves into a passive margin. In some cases, however, rifting stops before the continent splits in two; it becomes a low-lying trough that ﬁlls with sediment. Then, the rift remains as a permanent scar in the crust, deﬁned by a belt of faults, volcanic rocks, and a thick layer of sediment.

A major rift, known as the Basin and Range Province, breaks up the landscape of the western United States (figure above b). Here, movement on numerous faults tilted blocks of crust to form narrow mountain ranges, while sediment that eroded from the blocks ﬁlled the adjacent basins (the low areas between the ranges). Another active rift slices through eastern Africa; geoscientists aptly refer to it as the East African Rift (figure above c, d). To astronauts in orbit, the rift looks like a giant gash in the crust. On the ground, it consists of a deep trough bordered on both sides by high cliffs formed by faulting. Along the length of the rift, several major volcanoes smoke and fume; these include the snow-crested Mt. Kilimanjaro, towering over 6 km above the savannah. At its north end, the rift joins the Red Sea Ridge and the Gulf of Aden Ridge at a triple junction.

Collision

India was once a small, separate continent that lay far to the south of Asia. But subduction consumed the ocean between India and Asia, and India moved northward, ﬁnally slamming into the southern margin of Asia about 40 to 50 million years ago. Continental crust, unlike oceanic crust, is too buoyant to subduct. So when India collided with Asia, the attached oceanic plate broke off and sank down into the deep mantle while India pushed hard into and partly under Asia, squeezing the rocks and sediment that once lay between the two continents into the 8-km-high welt that we now know as the Himalayan Mountains. During this process, not only did the surface of the Earth rise, but the crust became thicker. The crust beneath a collisional mountain range can be up to 60 to 70 km thick, about twice the thickness of normal continental crust. The boundary between what was once two separate continents is called a suture; slivers of ocean crust may be trapped along a suture.

Continental collision (not to scale).

Geoscientists refer to the process during which two buoyant pieces of lithosphere converge and squeeze together as collision (figure above a, b). Some collisions involve two continents, whereas some involve continents and an island arc. When a collision is complete, the convergent plate boundary that once existed between the two colliding pieces ceases to exist. Collisions yield some of the most spectacular mountains on the planet, such as the Himalayas and the Alps. They also yielded major mountain ranges in the past, which subsequently eroded away so that today we see only their relicts. For example, the Appalachian Mountains in the eastern United States formed as a consequence of three collisions. After the last one, a collision between Africa and North America around 300 Ma, North America became part of the Pangaea supercontinent.

Credits: Stephen Marshak (Essentials of Geology)

9 Jun 2020

June 09, 2020 Comment

What Drives Plate Motion, and How Fast Do Plates Move?

We've now discussed the many facets of plate tectonics theory but to complete the story, we need to address a major question: “What drives plate motion?” When geoscientists ﬁrst proposed plate tectonics, they thought the process occurred simply because convective ﬂow in the asthenosphere actively dragged plates along, as if the plates were simply rafts on a ﬂowing river. Thus, early images depicting plate motion showed simple convection cells elliptical ﬂow paths in the asthenosphere. At ﬁrst glance, this hypothesis looked pretty good. But, on closer examination it became clear that a model of simple convection cells carrying plates on their backs can’t explain the complex geometry of plate boundaries and the great variety of plate motions that we observe on the Earth. Researchers now prefer a model in which convection, ridge push, and slab pull all contribute to driving plates. Let’s look at each of these phenomena in turn.

Convection is involved in plate motions in two ways. Recall that, at a mid-ocean ridge, hot asthenosphere rises and then cools to form oceanic lithosphere which slowly moves away from the ridge until, eventually, it sinks back into the mantle at a trench. Since the material forming the plate starts out hot, cools, and then sinks, we can view the plate itself as the top of a convection cell and plate motion as a form of convection. But in this view, convection is effectively a consequence of plate motion, not the cause. Can convection actually cause plates to move? The answer may come from studies which demonstrate that the interior of the mantle, beneath the plates, is indeed convecting on a very broad scale. Speciﬁcally, geologists have found that there are places where deeper, hotter asthenosphere is rising or upwelling, and places where shallower, colder asthenosphere is sinking or downwelling. Such asthenospheric ﬂow probably does exert a force on the base of plates. But the pattern of upwelling and downwelling on a global scale does not match the pattern of plate boundaries exactly. So, conceivably, asthenosphere-ﬂow may either speed up or slow down plates depending on the orientation of the ﬂow direction relative to the movement direction of the overlying plate.

Forces driving plate motions. Both ridge push and slab pull make plates move.

Ridge-push force develops simply because the lithosphere of mid-ocean ridges lies at a higher elevation than that of the adjacent abyssal plains (figure above a). To understand ridge-push force, imagine you have a glass containing a layer of water over a layer of honey. By tilting the glass momentarily and then returning it to its upright position, you can create a temporary slope in the boundary between these substances. While the boundary has this slope, gravity causes the weight of elevated honey to push against the glass adjacent to the side where the honey surface lies at lower elevation. The geometry of a midocean ridge resembles this situation, for sea ﬂoor of a midocean ridge is higher than sea ﬂoor of abyssal plains. Gravity causes the elevated lithosphere at the ridge axis to push on the lithosphere that lies farther from the axis, making it move away. As lithosphere moves away from the ridge axis, new hot asthenosphere rises to ﬁll the gap. Note that the local upward movement of asthenosphere beneath a mid-ocean ridge is a consequence of sea-ﬂoor spreading, not the cause.

Slab-pull force, the force that subducting, downgoing plates apply to oceanic lithosphere at a convergent margin, arises simply because lithosphere that was formed more than 10 million years ago is denser than asthenosphere, so it can sink into the asthenosphere (figure above b). Thus, once an oceanic plate starts to sink, it gradually pulls the rest of the plate along behind it, like an anchor pulling down the anchor line. This “pull” is the slab-pull force.

The Velocity of Plate Motions

Relative plate velocities: The blue arrows show the rate and direction at which the plate on one side of the boundary is moving with respect to the plate on the other side. The length of an arrow represents the velocity. Absolute plate velocities: The red arrows show the velocity of the plates with respect to a fixed point in the mantle.

How fast do plates move? It depends on your frame of reference. To illustrate this concept, imagine two cars speeding in the same direction down the highway. From the viewpoint of a tree along the side of the road, Car A zips by at 100 km an hour, while Car B moves at 80 km an hour. But relative to Car B, Car A moves at only 20 km an hour. Geologists use two different frames of reference for describing plate velocity. If we describe the movement of Plate A with respect to Plate B, then we are speaking about relative plate velocity. But if we describe the movement of both plates relative to a ﬁxed location in the mantle below the plates, then we are speaking of absolute plate velocity (figure above).

To determine relative plate motions, geoscientists measure the distance of a known magnetic anomaly from the axis of a mid-ocean ridge and then calculate the velocity of a plate relative to the ridge axis by applying this equation: plate velocity distance from the anomaly to the ridge axis divided by the age of the anomaly (velocity, by deﬁnition, is distance wtime). The velocity of the plate on one side of the ridge relative to the plate on the other is twice this value.

To estimate absolute plate motions, we can assume that the position of a mantle plume does not change much for a long time. If this is so, then the track of hot-spot volcanoes on the plate moving over the plume provides a record of the plate’s absolute velocity and indicates the direction of movement. (In reality, plumes are not completely ﬁxed; geologists use other, more complex methods to calculate absolute plate motions.)

Working from the calculations described above, geologists have determined that plate motions on Earth today occur at rates of about 1 to 15 cm per year about the rate that your ﬁngernails grow. But these rates, though small, can yield large displacements given the immensity of geologic time. At a rate of 10 cm per year, a plate can move 100 km in a million years! Can we detect such slow rates? Until the last decade, the answer was no. Now the answer is yes, because of satellites orbiting the Earth with global positioning system (GPS) technology. Automobile drivers use GPS receivers to ﬁnd their destinations, and geologists use them to monitor plate motions. If we calculate carefully enough, we can detect displacements of millimeters per year. In other words, we can now see the plates move this observation serves as the ultimate proof of plate tectonics.

Due to plate tectonics, the map of Earth‘s surface slowly changes. Here we see the assembly, and later the breakup, of Pangaea during the past 400 million years.

Taking into account many data sources that deﬁne the motion of plates, geologists have greatly reﬁned the image of continental drift that Wegener tried so hard to prove nearly a century ago. We can now see how the map of our planet’s surface has evolved radically during the past 400 million years (figure above), and even before.

Credits: Stephen Marshak (Essentials of Geology)

7 Jun 2020

June 07, 2020 Comment

Relation of Volcanism to Plate Tectonics

Different styles of volcanism occur at different locations on Earth. Most eruptions occur along plate boundaries, but major eruptions also occur at hot spots (figure above). We’ll now look at the settings in which eruptions occur, in the context of plate tectonics theory and see why different kinds of volcanoes form in different settings.

Mid-Ocean Ridges

Products of mid-ocean ridge volcanism cover 70% of our planet’s surface. We don’t generally see this volcanic activity, however, because the ocean hides most of it beneath a blanket of water. Mid-ocean ridge volcanoes, which develop along ﬁssures parallel to the ridge axis, are not all continuously active. Each one turns on and off in a time scale measured in tens to hundreds of years. They erupt basalt which, because it cools so quickly underwater, forms pillow-lava mounds. Water that heats up as it circulates through the crust near the magma chamber bursts out of hydrothermal (hot-water) vents along these mounds.

Convergent Boundaries

Most of the subaerial volcanoes on Earth lie along convergent plate boundaries (subduction zones). Subduction zones border over 60% of the Paciﬁc Ocean, creating a 20,000-km-long chain of volcanoes known as the Ring of Fire. Typically, individual volcanoes in volcanic arcs lie about 50 to 100 km apart. Some of these volcanoes grow on oceanic crust and become volcanic island arcs, such as the Marianas of the western Paciﬁc. Others grow on continental crust, building continental volcanic arcs such as the Cascade volcanic chain of Washington and Oregon or the Andes chain of South America.

Some of the lava that erupts at volcanic arcs consists of basalt derived by partial melting of the asthenosphere. At times when basaltic lava erupts, arc volcanoes display effusive activity. Some of the lava that erupts at arcs particularly at continental arcs is andesitic or rhyolitic. When these viscous lavas are erupting, arc volcanoes display explosive activity. Over the life of a volcano, effusive and explosive activity may alternate. As a result, the volcaniclastic debris erupted at one time may later be cloaked by a shell of hard lava, and thus will be protected from erosion. Eventually, arc volcanism can yield large stratovolcanoes, such as the elegant symmetrical cone of Mt. Fuji. Periods of growth, however, may be interrupted by explosions that leave behind a blasted-apart hulk like Mt. St. Helens (Volcanic Explosions to Remember below).

Due to the diversity of magmas that can form beneath rifts, rifts can host both basaltic ﬁssure eruptions, in which curtains of lava fountain up or linear chains of cinder cones develop, and explosive rhyolitic volcanoes. In some places, they even host stratovolcanoes such as Mt. Kilimanjaro in Africa.

Oceanic Hot-Spot Volcanoes

The interior of an oceanic hot-spot volcano is complicated. Initially, eruption produces pillow basalts. When the volcano emerges above sea level, it becomes a shield volcano. The margins of the island frequently undergo slumping, and the weight of the volcano pushes down the surface of the lithosphere. The Hawaiian islands exemplify this architecture.

When a hot-spot volcano ﬁrst forms on oceanic lithosphere, basaltic magma erupts at the surface of the sea ﬂoor. At ﬁrst, such submarine eruptions yield an irregular mound of pillow lava. With time, the volcano grows up above the sea surface and becomes an island. When the volcano emerges from the sea, the basalt lava that erupts no longer freezes so quickly, and thus ﬂows as a thin sheet over a great distance. Thousands of thin basalt ﬂows pile up, layer upon layer, to build a broad, dome-shaped shield volcano with gentle slopes (figure above). As the volcano grows, portions of it can’t resist the pull of gravity and slip seaward, creating large submarine slumps.

Continental Hot-Spot Volcanoes

Hot-spot volcanic activity in Yellowstone National Park.

Yellowstone National Park lies at the northeast end of a string of calderas, the oldest of which, at the southwest end of the track, erupted 16 million years ago (figure above a, b). Recent and ongoing activity beneath Yellowstone has yielded fascinating landforms, volcanic rock deposits, and geysers. Eruptions at the Yellowstone hot spot differ from those in Hawaii in an important way: unlike Hawaii, the Yellowstone hot spot erupts both basaltic lava and rhyolitic pyroclastic debris.

About 630,000 years ago, immense pyroclastic ﬂows, as well as convective clouds of ash and pumice lapilli, blasted out of the Yellowstone region. Close to the eruption, ignimbrites up to tens of meters thick formed, and ash and lapilli from the giant cloud sifted down over the United States as far east as the Mississippi River (figure above c). The eruption produced an immense caldera, up to 72 km across. When the debris settled, it blanketed an area of 2,500 square km with tuffs that, in the park, reached a thickness of 400 m. The park’s name reﬂects the brilliant colour of volcaniclastic debris (figure above d).

Flood-Basalt Eruptions

Flood basalt layers exposed on the wall of a canyon in Idaho.

In several locations around the world, huge sheets of low viscosity lava erupted out of ﬁssures and spread out in vast sheets. Geologists refer to the lava of these sheets as ﬂood basalt (figure above). Over time, many successive eruptions of ﬂood basalt can build up a broad plateau. The aggregate volume of rock in such a plateau may be so great (over 175,000 km3), that geologists also refer to the region as a large igneous province (LIP).

An example of a LIP, the Columbia River Plateau, occurs in Washington and Oregon. The basalt here, which erupted around 15 million years ago, reaches a thickness of 3.5 km. Geologists have identiﬁed about 300 individual ﬂows in the Columbia River Plateau. Lava in some of these ﬂows travelled great distances up to 600 km from its source. Eventually, basalt covered an area of 220,000 km. Even larger ﬂood-basalt provinces occur in eastern Siberia (an occurrence known as the Siberian Traps), the Deccan Plateau of India, the Paraná region of Brazil, and the Karroo Plateau of south Africa.

Iceland a Hot Spot on a Ridge

Iceland, a hot spot on the Mid-Atlantic Ridge.

Iceland is one of the few places on Earth where mid-ocean ridge volcanism has built a mound of basalt that protrudes above the sea. The island formed where a hot spot lies beneath the Mid-Atlantic Ridge the presence of this hot spot (probably due to an underlying mantle plume) means that far more magma erupted here than beneath other places along the ridge. Because Iceland straddles a divergent plate boundary, it is being stretched apart, with faults forming as a consequence. Indeed, the central part of the island is a narrow rift, in which the youngest volcanic rocks of the island have erupted (figure above a, b). This rift is the trace of the Mid-Atlantic Ridge. Faulting cracks the crust and so provides a conduit to a magma chamber. Thus, some eruptions on Iceland tend to be ﬁssure eruptions, yielding either curtains of lava that are many kilometres long or linear chains of small cinder cones.

Not all volcanic activity on Iceland occurs sub-aerially. Some eruptions take place under glaciers and melt large amounts of ice. When the melt water bursts through the edge of the glacier it becomes a devastating ﬂood called a jokulhlaup in Icelandic.

Volcanic Explosions to Remember

Explosions of volcanoes generate enduring images of destruction. The historical record shows a vast range in the volume of debris erupted, even though the largest observed eruption (Tambora in 1815) was small compared to a super-explosion that took place over 600,000 years ago in what is now Yellowstone National Park, Wyoming (figure above a). Let’s look at two notable examples of explosions.

Mt. St. Helens, a snow-crested stratovolcano in the Cascades of the north-western United States, had not erupted since 1857. However, geologic evidence suggested that the mountain had a violent past, punctuated by many explosive eruptions. On March 20, 1980, an earth quake announced that the volcano was awakening once again. A week later, a crater 80 m in diameter burst open at the summit and began emitting gas and pyroclastic debris. Geologists who set up monitoring stations to observe the volcano noted that its north side was beginning to bulge markedly, suggesting that the volcano was ﬁlling with magma, making the volcano expand like a balloon. Their concern that an eruption was imminent led local authorities to evacuate people in the area.

The climactic eruption came suddenly. At 8:32 A.M. on May 18, a geologist, David Johnston, monitoring the volcano from a distance of 10 km, shouted over his two-way radio, “Vancouver, Vancouver, this is it!” An earthquake had triggered a huge landslide that caused 3 cubic km of the volcano’s weakened north side to slide away. The sudden landslide released pressure on the magma in the volcano, causing a sudden and violent expansion of gases that blasted through the side of the volcano (figure above b). Rock, steam, and ash screamed north at the speed of sound and ﬂattened a forest and everything in it over an area of 600 square km (figure below c). Tragically, Johnston, along with 60 others, vanished forever.

Examples of explosive eruptions.

Seconds after the sideways blast, a vertical column carried about 540 million tons of ash (about 1 cubic km) 25 km into the sky, where the jet stream carried it away so that it was able to circle the globe. In towns near the volcano, a blizzard of ash choked roads and buried ﬁelds. Water saturated ash formed viscous slurries, or lahars, that ﬂooded river valleys, carrying away everything in their path.

When the eruption was ﬁnally over, the once cone-like peak of Mt. St. Helens had disappeared the summit now lay 440 m lower, and the once snow-covered mountain was a gray mound with a large gouge in one side. The volcano came alive again in 2004, but did not explode.

An even greater explosion happened in 1883. Krakatau, a volcano in the sea between Indonesia and Sumatra, where the Indian Ocean ﬂoor subducts beneath South-east Asia, had grown to become a 9-km-long island rising 800 m (2,600 ft) above the sea. On May 20, the island began to erupt with a series of large explosions, yielding ash that settled as far as 500 km away. Smaller explosions continued through June and July, and steam and ash rose from the island, forming a huge black cloud that rained ash into the surrounding straits. Ships sailing by couldn’t see where they were going, and their crews had to shovel ash off the decks.

Krakatau’s demise came at 10 A.M. on August 27, perhaps when the volcano cracked and the magma chamber suddenly ﬂooded with seawater. The resulting blast, ﬁve thousand times greater than the Hiroshima atomic bomb explosion, could be heard as far as 4,800 km away, and sub-audible sound waves travelled around the globe seven times. Giant waves pushed out by the explosion slammed into coastal towns, killing over 36,000 people. Near the volcano, a layer of ash up to 40 m thick accumulated. When the air ﬁnally cleared, Krakatau was gone, replaced by a submarine caldera some 300 m deep (figure above d). All told, the eruption shot 20 cubic km of rock into the sky. Some ash reached elevations of 27 km. Because of this ash, people around the world could view spectacular sunsets during the next several years.

Credits: Stephen Marshak (Essentials of Geology)

7 May 2020

May 07, 2020 Comment

3D Geological Model of Pakistan

We are glad to proportion you that our co-founder, Muhammad Qasim Mehmood with his magnificence fellows have organized a geological version which became presented at All Pakistan Science Fest hosted with the aid of UET Science Society at 20/04/17.

Here is the short introduction of model:

It is a 3D geological model of Pakistan that suggests specifically tectonic department of mountain degrees of Pakistan. The version demonstrates the predominant/well-known deposits of Pakistan like Petroleum, Minerals/Gemstones, Uranium, Coal along with different geographical functions like dams and rivers.

It is a non-operating version (length: five?6 feet approximately) supported especially via wooden forums and other cementing fabric. It is using thermocol sheets, maps, big paper sheets, graphs, paints and mechanical and medical goods as in keeping with the requirement of a specific version.

This unique version cover the topics of Plate Tectonics, Structural Geology, Economic Geology and others. Also the scholars have added the future components of Geo-economics like Kalabagh Dam, CPEC direction, oil and gasoline wells in Baluchistan and offshore wells in Arabian Sea near Gwadar.

The version is showing the following geological aspects of Pakistan:

1. Major Geological Basins of Pakistan i.E. Indus Basin and Balochistan Basin

2. Major Thrusts in Northern Pakistan

Main Karakoram Thrust (MKT)

Main Mantle Thrust (MMT)

Main Central Thrust (MCT)

Main Boundary Thrust (MBT)

Salt Range Thrust / Himalayan Frontal Thrust

3. Mountain Ranges of Pakistan

Some mountain ranges of Pakistan is shown on the model located in North-West to South-West of Pakistan which has important geological significance in distinguishing Indus Basin from Balochistan Basin

4. Famous Peaks of Pakistan

Mount Godwin-Austen (k2) - World's second maximum height

Nanga Parbat ( The Killer Mountain) - World's ninth highest height

Tirich Mir - highest height in Hindukush Range

five. Major Fuel of Pakistan

Oil wells in Potwar Plateau and in Balochistan

Gas wells in Sui, Balochistan - biggest gas reserve in Pakistan

Coal reserves in Thar - World's 16th biggest coal reserve in Pakistan

Uranium reserves in Siwalik Hills west of Dera Ghazi Khan

6. Famous Gem Stone of Pakistan

Emerald from Mingora, Swat

Aquamarine fromHunza Valley,Gilgit-Baltistan Tourmaline from Skardu District, Gilgit Baltistan

And two future potentialities for the improvement of Geo-economics of Pakistan:

1. Kalabagh Dam (to be made)

2. China Pakistan Economic Corridor -CPEC (construction under process) western route.

Following are some of the maps (received from internet) which we keep in mind all through the making ready of our model

Map showing Geological Basins of Pakistan

Source: GSP

Tectonic Map of Pakistan

Source: GSP

Political Map of Pakistan

Source: Unknown

Tectonic Map of Pakistan & India showing major regional thrusts

Source: Unknown

CPEC map

Source: CPEC website And some photos captured during the preparation of model

Cutting of thermocol sheet

coloring thermocol sheet with finishing paint

Hasan creating "finishing of paint" with paint spatula scraper

final look of Stage 1

Umer Amin sketching map on model and fixing sticks for projections and heights

sketches of mountain ranges and river tributaries and sticks for average height of each range and peak

all things are made perfect due to plotting of each point according to longitude and latitude

a great Atlas Book

maps and maps

a rough look of model showing mountain ranges made with Plaster of Paris

Completion of Stage 2

team work!!!

after using distemper paint

And finally after painting and drawing river tributaries, fixing sign boards of cities and much more, the model is:

3D Geological Model of Pakistan

Legend for the model

Geological tools, Gemstones, Rocks and Fuel (Coal and Crude Oil)

Featuring Qasim Mehmood (Co founder of Learning Geology), on left

and

Rana Faizan Saleem, my class mate and Founder of Geology for Beginners

Students of Institute of Geology, University of the Punjab

6 May 2020

May 06, 2020 Comment

Geologic Contacts

2. Intrusive contacts are those where one rock has intruded another

3. Fault contacts are those where rocks come into contact across fault zones.

Learn in element about fault right here

Following are the a few snap shots displaying each kind of geologic contact

Depositional Contacts

1. Angular Unconformity, Siccar Point, Scotland

This place is known as Siccar Point whichis the most important unconformity described by James Hutton (1726-1797) in support of his world-changing ideas on the origin and age of the Earth.

Here gently sloping strata of 370-million-year-old Famennian Late DevonianOld Red Sandstone and a basal layer of conglomerate overlie near vertical layers of 435-million-year-old lower Silurian Llandovery Epoch greywacke, with an interval of around 65 million years.

2. Cretaceous Sandstone overlying Conglomerate Kootenai Formation, SW Montana

Photo Courtesy: marlimillerphoto.Com

3. Dun Briste Sea Stack, IrelandDun Briste is a truly incredible site to see but must be visited to appreciate its splendour. It was once joined to the mainland. The sea stack stands 45 metres (150 feet) tall.

Dun Briste and the surrounding cliffs were formed around 350 million years ago (during the 'Lower Carboniferous Period'), when sea temperatures were much higher and the coastline at a greater distance away. There are many legends describing how the Sea Stack was formed but it is widely accepted that an arch leading to the rock collapsed during very rough sea conditions in 1393. This is remarkably recent in geological terms.