Relation of Volcanism to Plate Tectonics

Relation of Volcanism to Plate Tectonics

kacimo samy   3:28 ص

Relation of Volcanism to Plate Tectonics 

A map showing the distribution of volcanoes around the world and the basic geologic settings in which volcanoes form, in the contact of plate tectonics theory.

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).

Continental Rifts 

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)

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

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

kacimo samy   5:09 ص

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

Forces Acting on Plates 

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)

How Do Plate Boundaries Form and Die?

How Do Plate Boundaries Form and Die?

kacimo samy   4:57 ص

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)

Special Locations in the Plate Mosaic

Special Locations in the Plate Mosaic

kacimo samy   4:41 ص

Special Locations in the Plate Mosaic 

Triple Junctions 

Examples of triple junction. The triple junction are marked by dots.

Geologists refer to a place where three plate boundaries intersect as a triple junction, and name them after the types of boundaries that intersect. For example, the triple junction formed where the Southwest Indian Ocean Ridge intersects two arms of the Mid–Indian Ocean Ridge (this is the triple junction of the African, Antarctic, and Australian Plates) is a ridge-ridge-ridge triple junction (figure above a). The triple junction north of San  Francisco is a trench-transform-transform triple junction (figure above b).

Hot Spots 

 The dots represent the locations of selected hot-spot volcanoes. The red lines represent hot-spot tracks. The most recent volcano (dot) is at one end of this track. Some of these volcanoes are extinct, indicating that the mantle plume no longer exists. Some hot spots are fairly recent and do not have tracks. Dashed tracks were broken by sea-floor spreading.

Most subaerial (above sea level) volcanoes are situated in the volcanic arcs that border trenches. Volcanoes also lie along mid-ocean ridges, but ocean water hides most of them. The volcanoes of volcanic arcs and mid-ocean ridges are plate boundary volcanoes, in that they formed as a consequence of movement along the boundary. Not all volcanoes on Earth are plate-boundary volcanoes, however. Worldwide, geoscientists have identified about 100 volcanoes that exist as isolated points and are not a consequence of movement at a plate boundary. These are called hotspot volcanoes, or simply hot spots (figure above). Most hot spots are located in the interiors of plates, away from the boundaries, but a few lie along mid-ocean ridges. What causes hot-spot volcanoes? In the early 1960s, J. Tuzo Wilson noted that active hot-spot volcanoes (examples that are erupting or may erupt in the future) occur at the end of a chain of dead volcanic islands and seamounts (formerly active volcanoes that will never erupt again). This configuration is different from that of volcanic arcs along convergent plate boundaries at volcanic arcs, all of the volcanoes are active. With this image in mind, Wilson suggested that the position of the heat source causing a hotspot volcano is fixed, relative to the moving plate. In Wilson’s model, the active volcano represents the present-day location of the heat source, whereas the chain of dead volcanic islands represents locations on the plate that were once over the heat source but progressively moved off.

The deep mantle plume hypothesis for the formation of hot-spot tracks.

A few years later, researchers suggested that the heat source for hot spots is a mantle plume, a column of very hot rock rising up through the mantle to the base of the lithosphere (figure above a–d). In this model, plumes originate deep in the mantle. Rock in the plume, though solid, is soft enough to flow, and rises buoyantly because it is less dense  than surrounding cooler rock. When the hot rock of the plume reaches the base of the lithosphere, it partially melts and produces magma that seeps up through the lithosphere to the Earth’s surface. The chain of extinct volcanoes, or hot-spot track, forms when the overlying plate moves over a fixed plume. This movement slowly carries the volcano off the top of the plume, so that it becomes extinct. A new, younger volcano grows over the plume. 

The Hawaiian chain provides an example of the volcanism associated with a hot-spot track. Volcanic  eruptions occur today only on the big island of Hawaii. Other islands to the northwest are remnants of dead volcanoes, the oldest of which is Kauai. To the northwest of Kauai, still older volcanic remnants are found. About 1,750 km northwest of Midway Island, the track bends in a more northerly direction, and the volcanic remnants no longer poke above sea level; we refer to this northerly trending segment as the Emperor seamount chain. Geologists suggest that the bend is due to a change in the direction of Pacific Plate motion at about 40 Ma. 

Some hot spots lie within continents. For example, several have been active in the interior of Africa, and one now underlies Yellowstone National Park. The famous geysers (natural steam and hot-water fountains) of Yellowstone exist because hot magma, formed above the Yellowstone hot spot, lies not far below the surface of the park. While most hot spots, such as Hawaii and Yellowstone, occur in the interior of plates, away from plate boundaries, a few are positioned at points on mid-ocean ridges. The additional magma production associated with such hot spots causes a portion of the ridge to grow into a mound that can rise significantly above normal ridgeaxis depths and protrude above the sea surface. Iceland, for example, is the product of hot-spot volcanism on the axis of the Mid-Atlantic Ridge.
Credits: Stephen Marshak (Essentials of Geology)

Transform Plate Boundaries

Transform Plate Boundaries

kacimo samy   3:02 ص

Transform Plate Boundaries

The concept of transform faulting.

When researchers began to explore the bathymetry of midocean ridges in detail, they discovered that mid-ocean ridges are not long, uninterrupted lines, but rather consist of short segments that appear to be offset laterally from each other (figure above a) by narrow belts of broken and irregular sea floor. These belts, or fracture zones, lie roughly at right angles to the ridge segments, intersect the ends of the segments, and extend beyond the ends of the segments. Originally, researchers incorrectly assumed that the entire length of each fracture zone was a fault, and that slip on a fracture zone had displaced segments of the mid-ocean ridge sideways, relative to each other. In other words, they imagined that a mid-ocean ridge initiated as a continuous, fence-like line that only later was broken up by faulting. But when information about the distribution of earthquakes along mid-ocean ridges became available, it was clear that this model could not be correct. Earthquakes, and therefore active fault slip, occur only on the segment of a fracture zone that lies between two ridge segments. The portions of fracture zones that extend beyond the edges of ridge segments, out into the abyssal plain, are not 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 introduced the term transform boundary, or transform fault, for the actively slipping segment of a fracture zone between two ridge segments, and he pointed out that these are a third type of plate boundary. At a transform boundary, one plate slides sideways past another, but no new plate forms and no old plate is consumed. Transform boundaries are, therefore, defined by a vertical fault on which the slip direction parallels the Earth’s surface. The slip breaks up the crust and forms 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)

Convergent Plate Boundaries and Subduction

Convergent Plate Boundaries and Subduction

kacimo samy   2:56 ص

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)

Divergent Plate Boundaries and Sea-Floor Spreading

Divergent Plate Boundaries and Sea-Floor Spreading

kacimo samy   2:49 ص

Divergent Plate Boundaries and Sea-Floor Spreading 

The process of sea-floor spreading.

At a divergent boundary, or spreading boundary, two oceanic plates move apart by the process of sea-floor spreading. Note that an open space does not develop between diverging plates. Rather, as the plates move apart, new oceanic lithosphere forms continually along the divergent boundary (figure above a). This process takes place at a submarine mountain range called a mid-ocean ridge that rises 2 km above the adjacent abyssal plains of the ocean. Thus, geologists commonly refer to a divergent boundary as a mid-ocean ridge, or simply a ridge. Water depth above ridges averages about 2.5  km. 

To characterize a divergent boundary more completely, let’s look at one mid-ocean ridge in more detail (figure above b). The Mid-Atlantic Ridge extends from the waters between northern  Greenland and northern Scandinavia southward across the equator to the latitude of the southern tip of South America. Geologists have found that the formation of new sea floor takes place only along the axis (centerline) of the ridge, which is marked by an elongate valley. The sea floor slopes away, reaching the depth of the abyssal plain (4 to 5 km) at a distance of about 500 to 800 km from the ridge axis. Roughly speaking, the Mid-Atlantic Ridge is symmetrical its eastern half looks like a mirror image of its western half. The ridge consists, along its length, of short segments (tens to hundreds of km long) that step over at breaks that, as we noted earlier, are called fracture zones.

How Does Oceanic Crust Form  at a Mid-Ocean Ridge? 

As sea-floor spreading takes place, hot asthenosphere rises beneath the ridge and begins to melt, and molten rock, or magma, forms (figure above c). Magma has a lower density than solid rock, so it behaves buoyantly and rises, as oil rises above vinegar in salad dressing. Molten rock eventually accumulates in the crust below the ridge axis, filling a region called a magma chamber. As the magma cools, it turns into a mush of crystals. Some of the magma solidifies completely along the side of the chamber to make the coarse-grained, mafic igneous rock called gabbro. The rest rises still higher to fill vertical cracks, where it solidifies and forms wall-like sheets, or dikes, of basalt. Some magma rises all the way to the surface of the sea floor at the ridge axis and spills out of small submarine volcanoes. The resulting lava cools to form a layer of basalt blobs called pillows. Observers in research submarines have detected chimneys spewing hot, mineralized water rising from cracks in the sea floor along the ridge axis. These chimneys are called black smokers because the water they emit looks like a cloud of dark smoke; the colour comes from a suspension of tiny mineral grains that precipitate in the water the instant that the water cools (figure below). 

 A column of superhot water gushing from a vent known as a black smoker along the mid-ocean ridge. A local ecosystem of bacteria, shrimp, and worms lives around the vent.

As soon as it forms, new oceanic crust moves away from the ridge axis, and when this happens, more magma rises from below, so still more crust forms. In other words, like a vast, continuously moving conveyor belt, magma from the mantle rises to the Earth’s surface at the ridge, solidifies to form oceanic crust, and then moves laterally away from the ridge. Because all sea floor forms at mid-ocean ridges, the youngest sea floor occurs on either side of the ridge axis, and sea floor becomes progressively older away from the ridge. In the Atlantic Ocean, the oldest sea floor, therefore, lies adjacent to the passive continental margins on either side of the ocean (figure below). The oldest ocean floor on our planet underlies the western Pacific Ocean; this crust formed about 200 million years ago.

This map of the world shows the age of the sea floor. Note how the sea floor grows older with increasing distance from the ridge axis. (Ma = million years ago.)

The tension (stretching force) applied to newly formed solid crust as spreading takes place breaks the crust, resulting in the formation of faults. Slip on the faults causes divergent boundary earthquakes and produces numerous cliffs, or scarps, that lie parallel to the ridge axis.

How Does the Lithospheric Mantle  Form at a Mid-Ocean Ridge? 

So far, we've seen how oceanic crust forms at mid-ocean ridges. How does the mantle part of the oceanic lithosphere form? This part consists of the cooler uppermost layer of the mantle, in which temperatures are less than about 1,280°C. At the ridge axis, such temperatures occur almost at the base of the crust, because of the presence of rising hot asthenosphere and hot magma, so lithospheric mantle beneath the ridge axis effectively doesn’t exist. But as the newly formed oceanic crust moves away from the ridge axis, the crust and the uppermost mantle directly beneath it gradually cool by losing heat to the ocean above. As soon as mantle rock cools below 1,280°C, it becomes, by definition, part of the lithosphere.

Changes accompanying the aging of lithosphere.

As oceanic lithosphere continues to move away from the ridge axis, it continues to cool, so the lithospheric mantle, and therefore the oceanic lithosphere as a whole, grows progressively thicker (figure above a, b). Note that this process doesn’t change the thickness of the oceanic crust, for the crust formed entirely at the ridge axis. The rate at which cooling and lithospheric thickening occur decreases progressively with increasing  distance from the ridge axis. In fact, by the time the lithosphere is about 80 million years old, it has just about reached its maximum thickness. As lithosphere thickens and gets cooler and denser, it sinks down into the asthenosphere, like a ship taking on ballast. Thus, the ocean is deeper over older ocean floor than over younger ocean floor.

Credits: Stephen Marshak (Essentials of Geology)

What Do We Mean by Plate Tectonics?

What Do We Mean by Plate Tectonics?

kacimo samy   2:32 ص

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.

The Concept of a  Lithosphere Plate 

 Nature of the lithosphere and its behaviour.

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 location 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 background provided above, we can restate plate tectonics theory concisely as follows. The Earth’s lithosphere is divided into plates that move relative to each other. As a plate moves, its internal area remains mostly, but not perfectly, rigid and intact. But rock along plate boundaries undergoes intense deformation (cracking, sliding, bending, stretching, and squashing) as the plate grinds or scrapes against its neighbours or pulls away from its neighbours. As plates move, so do the continents that form part of the plates. Because of plate tectonics, the map of Earth’s surface constantly changes.

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 types of plate boundaries differ based on the nature of relative movement.

Geologists define three types of plate boundaries, based simply on the relative motions of the plates on either side of the boundary (figure above a–c). A boundary at which two plates move apart from each other is a divergent boundary. A boundary at which two plates move toward each other so that one plate sinks beneath the other is a convergent boundary. And a boundary at which two plates slide sideways past each other is a transform boundary.

Credits: Stephen Marshak (Essentials of Geology)

Evidence for Sea-Floor Spreading

Evidence for Sea-Floor Spreading

kacimo samy   8:48 ص

Evidence for Sea-Floor Spreading 

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.

Marine Magnetic Anomalies

Recognizing anomalies 

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.

Magnetic reversals

Magnetic polarity reversals and the chronology of 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)

The Discovery of Sea-Floor Spreading

The Discovery of Sea-Floor Spreading

kacimo samy   7:43 ص

The Discovery of Sea-Floor Spreading

New Images of Sea-Floor Bathymetry 

Bathymetry of mid-ocean ridges and abyssal plains.

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 floor.

• 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 discovered many important characteristics of the sea-floor crust. These discoveries led them to realize that oceanic crust differs from continental crust, and that bathymetric features of the ocean floor provide clues to the origin of the crust. Specifically: 

• 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 showing the distribution of earthquake locations in the ocean basins. Note that earthquakes occur in belts.

Harry Hess and His “Essay in Geopoetry” 

Harry Hess's basic concept of sea-floor spreading. Hess implied, incorrectly, that only the crust moved. We will see that this sketch 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 sea-floor-spreading hypothesis instantly provided the long sought explanation of how continental “drift” occurs. Continents passively move apart as the sea floor between them spreads at mid-ocean ridges, and they passively move together as the sea floor between them sinks back into the mantle at trenches. (As we will see later, geologists now realize that it is the lithosphere that moves, not just the crust.) Thus, sea-floor spreading proved to be an important step on the route to plate tectonics the idea seemed so good that Hess referred to his description of it as “an essay in geopoetry.” But first, the idea needed to be tested, and other key discoveries would have to take place before the whole theory of plate tectonics could come together.