Seismic Study of Earth’s Interior

Seismic Study of Earth’s Interior

kacimo samy   2:08 م

Seismic Study of Earth’s Interior 

Let’s now utilize your knowledge of seismic velocity, refraction, and reflection to see how each of the major layer boundaries inside the Earth was discovered.

Discovering the Crust-Mantle Boundary 

Discovery of the Moho.

The concept that seismic waves refract at boundaries between different layers led to the first documentation of the core-mantle boundary. In 1909, Andrija Mohorovicic, a Croatian seismologist, noted that P-waves arriving at seismometer stations less than 200 km from the epicentre travelled at an average speed of 6 km per second, whereas P-waves arriving at seismometers more than 200 km from the epicentre travelled at an average speed of 8 km per second. To explain this observation, he suggested that P-waves reaching nearby seismometers followed a shallow path through the crust, in which they travelled relatively slowly, whereas P-waves reaching distant seismometers followed a deeper path through the mantle, in which they travelled relatively rapidly (figure above a, b).

Calculations based on this observation require the crust mantle boundary beneath most continental regions to be at a depth of about 35 to 40 km. Later studies showed that the crust-mantle boundary beneath oceanic regions lies at a depth of about 7 to 10 km. The crust-mantle boundary is now called the Moho, in honour of Mohorovicic.

Defining the Structure of the Mantle 

The velocity of P-waves in the mantle changes because the physical properties of the mantle change with depth.

By studying travel times, seismologists have determined that seismic waves travel at different speeds at different depths in the mantle. Between a depth of about 100 and 200 km in the asthenosphere beneath oceanic lithosphere, seismic velocities are slower than in the overlying lithospheric mantle (figure above a). This 100- to 200-km-deep layer is called the low-velocity zone (LVZ); here, the prevailing temperature and pressure conditions cause peridotite to partially melt by up to 2%. The melt, a liquid, coats solid grains and fills voids between grains. Because seismic waves travel more slowly through liquids than through solids, the coatings of melt slow seismic waves down. In the context of plate tectonics theory, the low-velocity zone is the weak layer on which oceanic lithosphere plates move. Below the low-velocity zone, the asthenosphere does not contain melt. (Of note, seismologists do not find a well-developed low-velocity zone beneath continents.) 

Below about 200 km, seismic-wave velocities in the mantle increase with depth (figure above b). Seismologists interpret this increase to mean that mantle peridotite becomes progressively less compressible, more rigid, and denser with depth. This proposal makes sense, considering that the weight of overlying rock increases with depth, and as pressure increases, the atoms making up rock squeeze together more tightly and are not so free to move. Because of refraction, the increase in seismic  velocity with depth causes seismic rays to curve in the mantle. To understand the shape of a curved ray, let’s represent a portion of the mantle by a series of imaginary layers, each of which has a slightly greater seismic-wave velocity than the layer above (figure above c). Every time a seismic ray crosses the boundary between adjacent layers, it refracts a little toward the boundary. After the ray has crossed several layers, it has bent so much that it begins to head back up toward the top of the stack. Now if we replace the stack of distinct layers with a single layer in which velocity increases with depth at a constant rate, the wave follows a smoothly curving path (figure above d). 

At depths between 410 km and 660 km, seismic velocity increases in a series of abrupt steps (figure above a), so the stack of layers in figure above c is actually a somewhat realistic image. A major step occurs at a depth of 660 km. Experiments suggest that such seismic-velocity discontinuities occur at depths where pressure abruptly causes atoms in minerals to rearrange and pack together more tightly, thereby changing the rock’s compressibility and rigidity. Researchers are still unsure if chemical changes also occur at the discontinuities. Because of these seismic-velocity discontinuities, now can see more clearly, seismologists subdivide the mantle into the upper mantle (above 660 km), the transition zone (between 410 and 660 km), and lower mantle (below 660 km). Note that seismologists consider the transition zone to be part of the upper mantle. 

Discovering the Core-Mantle Boundary 

Shadow zones and the discovery of the Earth's core. 

During the first decade of the 20th century, seismologists installed seismometers at many stations around the world, expecting to be able to record waves produced by a large earthquake anywhere on Earth. In 1914, one of these seismologists, Beno Gutenberg, discovered that P-waves from a given earthquake do not arrive at seismometers lying in a band at a distance of between 103° and 143°, as measured along the surface of the Earth from the earthquake epicentre. This band is now called the P-wave shadow zone (figure above a). If the density of the Earth increased gradually with depth all the way to the centre, the shadow zone would not exist because rays passing into the interior would curve up and reach every point on the surface. Thus, the presence of a shadow zone means that deep in the Earth a major interface exists where seismic waves abruptly refract down (implying that the velocity of seismic waves suddenly decreases). This interface, the core-mantle boundary, lies at a depth of about 2,900 km. 

To see why the P-wave shadow zone exists, follow the two seismic rays labelled A and B in figure above a. Ray A curves smoothly in the mantle (we are ignoring seismic-velocity discontinuities in the mantle) and passes just above the core mantle boundary before returning to the surface. It reaches the surface 103° from the epicentre. In contrast, Ray B penetrates the boundary and refracts down into the core. Ray B then curves through the core and refracts again when it crosses back into the mantle. As a consequence, Ray B intersects the surface at more than 143° from the epicentre.

Discovering the Nature of the Core 

Based on the study of meteorites thought to be fragments of a large planetesimal’s interior, and on density calculations, seismologists concluded that the core consists of iron alloy. This alloy contains about 85% iron, 5% nickel, and 10% of a lighter element (probably oxygen, silicon, and/or sulphur). But the downward bending of seismic waves when they pass from the mantle down into the core indicates that seismic velocities in, at least, the outer core are slower than in the mantle. Thus, even though the core is deeper and denser than the mantle, at least the outer part of the core must be less rigid than the mantle. How can this be?

Seismologists found that S-waves do not arrive at stations located between 103° and 180° from the epicentre (a band called the S-wave shadow zone). This means that S-waves cannot pass through the core at all otherwise, an S-wave headed straight down through the Earth would appear on the other side. S-waves are shear waves, which by their nature can travel only through solids. Thus, the fact that S-waves do not pass through the core means that the core, or at least part of it, consists of liquid (figure above b). 

At first, seismologists thought that the entire core might be liquid iron alloy. But in 1936, a Danish seismologist, Inge Lehmann, discovered that P-waves passing through the core reflected off a boundary within the core. She proposed that the core is made up of two parts: an outer core consisting of liquid iron alloy and an inner core consisting of solid iron alloy. Lehmann’s work defined the existence of the inner core but could not locate the depth at which the inner core-outer core interface occurs. This depth was eventually located by measuring the exact time it took for seismic waves to penetrate the Earth, bounce off the inner core–outer core boundary, and return to the surface (figure above c). The measurements showed that the inner core–outer core boundary occurs at a depth of about 5,155 km. 

Why does the core have two layers a liquid outer layer and a solid inner one? An examination of figure above d provides some insight. This graph shows two curves: (1) the geotherm, which indicates how temperature changes with increasing depth in the Earth; and (2) the melting curve, which indicates that the temperature at which materials melt changes with increasing depth in the Earth. At temperatures to the right of the melting curve, Earth materials are molten. As the graph shows, the geotherm lies to the left of the melting curve through most of the mantle and in the inner core. This means that temperatures in most of the mantle and in the inner core are not high enough to cause melting, under the very high pressures found in these regions, so these regions are solid. But the geotherm lies to the right of the melting curve in the low-velocity zone of the mantle and in the outer core, so these regions contain molten material. 

A Modern Image of Earth’s Layers 

The velocity-versus-depth profile of the whole Earth.

Through painstaking effort, seismologists used data from this array to develop a graph, known as a velocity-versus depth curve, that shows the depths at which seismic velocity suddenly changes. These changes define the principal layers and sub-layers in the Earth (figure above). Note that the graph does not show a velocity for S-waves in the outer core because S-waves cannot travel through molten iron (a liquid). 

Seismic Tomography 

In recent years, seismologists have developed a technique, called seismic tomography, to produce three-dimensional images of variation in seismic velocities in the Earth’s interior. This technique resembles the method used to produce three-dimensional CAT (or CT) scans of the human body. In seismic tomography studies, researchers compare the observed travel time of seismic waves following a specific ray path with the predicted travel time that waves following the same path would have if the average velocity-versus-depth model depicted by Figures 2 and and 3 were completely correct. 

Tomographic images of the Earth’s interior, and their interpretation.

Tomographic studies emphasize that the simple onion-like layered image of the Earth, with velocities increasing with depth at the same rate everywhere, is an oversimplification. In reality, the velocities of seismic waves vary significantly with location at a given depth. Results of tomographic studies can be displayed by three-dimensional models, cross sections, or maps (figure above a,  b). Generally, warmer colours (reds) on these images indicate slower and presumably warmer regions, whereas cooler colours (blues and purples) indicate faster and presumably cooler regions.

Even though many important questions remain, tomography has led geologists to picture the Earth’s insides as a dynamic place (figure above c). This image should become even clearer, because a major research initiative, called Earth Scope, has begun. This initiative involves placing hundreds of seismometers in an array across the United States. Just as digital photographs have higher resolution when taken by a camera with a 14-mega-pixel sensor than one with a 2-mega-pixel sensor, the greater number of seismometers in the Earth Scope array provides a higher-resolution tomographic image of the Earth’s interior. 

Seismic-Reflection Profiling 

Seismic reflection profiling. The method allows geologists to see underground.

Seismic techniques are also letting us fine-tune our image of the upper crust. During the past half century, geologists have found that by exploding dynamite, by banging large weights against the Earth’s surface, or by releasing bursts of compressed air into the water, they can create artificial seismic waves that propagate down into the Earth and reflect off the boundaries between different layers of rock in the crust. By recording the time at which these reflected waves return to the surface, geologists can determine the depth to these boundaries. With this information, they can produce a cross-sectional view of the crust called a seismic-reflection profile (figure above a, b). This image can define subsurface bedding and stratigraphic formation contacts, and can reveal the presence of subsurface folds (bends in layers) and faults. Oil companies obtain many seismic-reflection profiles, despite their high cost, because they allow geologists to identify likely locations for oil and gas reserves underground.

In the past decade, computers have become so sophisticated that geologists can now produce three-dimensional seismic-reflection images of the crust. From the 3-D data, they can produce profiles and maps in any orientation desired. These provide so much detail that geologists can even trace out a ribbon of sand, representing the channel of an ancient stream, now buried kilometres below the Earth’s surface.

Credits: Stephen Marshak (Essentials of Geology)

The Movement of Seismic Waves Through the Earth

The Movement of Seismic Waves Through the Earth

kacimo samy   1:32 م

The Movement of Seismic Waves Through the Earth 

Wave Fronts and Travel Times 

The propagation of earthquake waves.

The energy released by an earthquake moves through rock in the form of waves, just as waves propagate outward from the impact point of a pebble on the surface of a pond. The boundary between the rock through which a wave has passed and the rock through which it has not yet passed is called a wave front. In 3-D, a wave front expands outward from the earthquake focus like a growing bubble. We can represent a succession of waves in a drawing by a series of concentric wave fronts. The changing position of an imaginary point on a wave front as the front moves through rock is called a seismic ray. You can picture a seismic ray as a line drawn perpendicular to a wave front; each point on a curving wave front follows a slightly different ray (figure above a). The time it takes for a wave to travel from the focus to a seismometer along a given ray is the travel time along that ray.
 

The ability of a seismic wave to travel through a certain material, as well as the velocity at which it travels, depends on the character of the material. Factors such as density (mass per unit volume), rigidity (how stiff or resistant to bending a material is), and compressibility (how easily a material’s volume changes in response to squashing) all affect seismic wave movement. Studies of seismic waves reveal the following:

• Seismic waves travel at different velocities in different rock types (figure above b). For example, P-waves travel at 8 km per second in peridotite (an ultramafic igneous rock), but at only 3.5 km per second in sandstone (a porous sedimentary rock). Therefore, waves accelerate or slow down if they pass from one rock type into another. 
• Seismic waves travel more slowly in magma than in solid rock of the same composition, and more slowly in molten iron alloy than in solid iron alloy (figure above c). 
• Both P-waves and S-waves can travel through a solid, but only P-waves can travel through a liquid (figure above d).

Reflection and Refraction of Wave Energy 

Refraction and reflection of waves.

Shine a flash light into a container of water so that the light ray hits the boundary (or interface) between water and air at an angle. Some of the light bounces off the water surface and heads back up into the air, while some enters the water  (figure above a). The light ray that enters the water bends at the air water boundary so the angle between the ray and the boundary in the air is different from the angle between the ray and the boundary in the water. Physicists refer to the light ray that bounces off the air-water boundary and heads back into the air as the reflected ray, and the ray that bends at the boundary as the refracted ray. The phenomenon of bouncing off is reflection, and the phenomenon of bending is refraction. Wave reflection and refraction take place at the interface between two materials, if the wave travels at different velocities in the two materials.

Seismic energy travels in the form of waves, so seismic waves reflect and/or refract when reaching the interface between two rock layers if the waves travel at different velocities in the two layers. For example, imagine a layer of sandstone overlying a layer of basalt. Seismic velocities in sandstone are slower than in basalt, so as seismic waves reach the boundary,  some reflect and some refract. The amount and direction of refraction at a boundary depend on the contrast in wave velocity across the boundary and on the angle at which a wave hits the interface. As a rule, if waves enter a material through which they will travel more slowly, the rays representing the waves bend down and away from the interface. For example, the light ray in figure above b bends down when hitting the air-water boundary because light travels more slowly in water. This relation makes sense if you picture a car driving from a paved surface diagonally onto a sandy beach the wheel that rolls onto the sand first slows down relative to the wheel still on the pavement, causing the car to turn toward the sand. Alternatively, if the ray were to pass from a layer in which it travels slowly into one in which it travels more rapidly, the rays representing the waves would bend up and toward the interface (figure above b).

Credits: Stephen Marshak (Essentials of geology)

What is Earth made of?

What is Earth made of?

kacimo samy   6:31 ص

What is Earth made of?

The proportions of major elements making up the mass of the whole Earth.

At this point, we leave our fantasy space voyage and turn our attention inward to the materials that make up the solid Earth, because we need to be aware of these before we can discuss the architecture of the Earth’s interior. Let’s begin by reiterating that the Earth consists mostly of elements produced by fusion reactions in stars and by supernova explosions. Only four elements (iron, oxygen, silicon, and magnesium) make up 91.2% of the Earth’s mass; the remaining 8.8% consists of the other 88 elements (figure above). The elements of the Earth comprise a great variety of materials.

• Organic chemicals. Carbon-containing compounds that either occur in living organisms or have characteristics that resemble compounds in living organisms are called organic chemicals. 
• Minerals. A solid, natural substance in which atoms are arranged in an orderly pattern is a mineral. A single coherent sample of a mineral that grew to its present shape is a crystal, whereas an irregularly shaped sample, or a fragment derived from a once-larger crystal or cluster of crystals, is a grain. 
• Glasses. A solid in which atoms are not arranged in an orderly pattern is called glass. 
• Rocks. Aggregates of mineral crystals or grains, or masses of natural glass, are called rocks. Geologists recognize three main groups of rocks. (1) Igneous rocks develop when hot molten (liquid) rock cools and freezes solid. (2) Sedimentary rocks form from grains that break off pre-existing rock and become cemented together, or from minerals that precipitate out of a water solution. (3) Metamorphic rocks form when pre-existing rocks change in response to heat and pressure. 
• Sediment. An accumulation of loose mineral grains (grains that have not stuck together) is called sediment. 
• Metals. A solid composed of metal atoms (such as iron, aluminium, copper, and tin) is called a metal. An alloy is a mixture containing more than one type of metal atom. 
• Melts. A melt forms when solid materials become hot and transform into liquid. Molten rock is a type of melt geologists distinguish between magma, which is molten rock beneath the Earth’s surface, and lava, molten rock that has flowed out onto the Earth’s surface. 
• Volatiles. Materials that easily transform into gas at the relatively low temperatures found at the Earth’s surface are called volatiles. 

The most common minerals in the Earth contain silica (a compound of silicon and oxygen) mixed in varying proportions with other elements. These minerals are called silicate minerals. Not surprisingly, rocks composed of silicate minerals are silicate rocks. Geologists distinguish four classes of igneous silicate rocks based, in essence, on the proportion of silica to iron and magnesium. In order, from greatest to least proportion of silica to iron and magnesium, these classes are felsic (or silicic), intermediate, mafic, and ultramafic. As the proportion of silica in a rock increases, the density (mass per unit volume) decreases. Thus, felsic rocks are less dense than mafic rocks. Many different rock types occur in each class, as will be discussed in detail in Chapters 4 through 7. For now, we introduce the four rock types whose names we need to know for our discussion of the Earth’s layers that follows. These are (1) granite, a felsic rock with large grains; (2) gabbro, a mafic rock with large grains; (3) basalt, a mafic rock with small grains; and (4) peridotite, an ultramafic rock with large grains.

Discovering the Earth’s Internal Layers

People have speculated about what’s inside our planet since ancient times. What is the source of incandescent lavas that spew from volcanoes, of precious gems and metals found in mines, of sparkling mineral waters that bubble from springs, and of the mysterious forces that shake the ground and topple buildings? In ancient Greece and Rome, the subsurface was the underworld, Hades, home of the dead, a region of fire and sulphurous fumes. Perhaps this image was inspired by the molten rock and smoke emitted by the volcanoes of the Mediterranean region. In the 18th and 19th centuries, European writers thought the Earth’s interior resembled a sponge, containing open caverns variously filled with molten rock, water, or air. In fact, in the popular 1864 novel Journey to the Centre of the Earth, by the French author Jules Verne, three explorers hike through interconnected caverns down to the Earth’s centre.
How can we explore the interior for real? We can’t dig or drill down very far. Indeed, the deepest mine penetrates only about 3.5 km beneath the surface of South Africa. And the deepest drill hole probes only 12 km below the surface of northern Russia compared with the 6,371 km radius of the Earth, this hole makes it less than 0.2% of the way to the centre and is nothing more than a pinprick. Our modern image of the Earth’s interior, one made up of distinct layers, is the end product of many discoveries made during the past 200 years.
The first clue that led away from Jules Verne’s sponge image came when researchers successfully measured the mass of the whole Earth, and from this information derived its average density. They found that the average density of our planet far exceeds the density of common rocks found on the surface. Thus, the interior of the Earth must contain denser material than its outermost layer and can’t possibly be full of holes. In fact, the mass of the Earth overall is so great that the planet must contain a large amount of metal. Since the Earth is close to being a sphere, the metal must be concentrated near the centre. Otherwise, centrifugal force due to the spin of the Earth on its axis would pull the equator out, and the planet would become a disk. (To picture why, consider that when you swing a hammer, your hand feels more force if you hold the end of the light wooden shaft, rather than the heavy metal head.) Finally, researchers realized that, though molten rock occasionally oozes out of the interior at volcanoes, the interior must be mostly solid, because if it weren't, the land surface would rise and fall due to tidal forces much more than it does.

An early image of Earth’s internal layers.

Eventually, researchers concluded that the Earth resembled a hard-boiled egg, in that it had three principal layers: a not-so-dense crust (like an eggshell, composed of rocks such as granite, basalt, and gabbro), a denser solid mantle in the middle (the “white,” composed of a then-unknown material), and a very dense core (the “yolk,” composed of an unknown metal) (figure above). Clearly, many questions remained. How thick are the layers? Are the boundaries between layers sharp or gradational? And what exactly are the layers composed of?

Clues from the Study of Earthquakes: Refining the Image

Faulting and earthquakes.

When rock within the outer portion of the Earth suddenly breaks and slips along a fracture called a fault, it generates shock waves (abrupt vibrations), called seismic waves, that travel through the surrounding rock outward from the break. Where these waves cause the surface of the Earth to vibrate, people feel an earthquake, an episode of ground shaking. You can simulate this process, at a small scale, when you break a stick between your hands and feel the snap with your hands (figure above).

In the late 19th century, geologists learned that earthquake energy could travel, in the form of waves, all the way through the Earth’s interior from one side to the other. Geologists immediately realized that the study of earthquake waves travelling through the Earth might provide a tool for exploring the Earth’s insides, much as ultrasound today helps doctors study a patient’s insides. Specifically, laboratory measurements demonstrated that earthquake waves travel at different velocities (speeds) through different materials. Thus, by detecting depths at which velocities suddenly change, geoscientists pinpointed the boundaries between layers and even recognized subtler boundaries within layers. For example, such studies led geoscientists to subdivide the mantle into the upper mantle and lower mantle, and subdivide the core into the inner core and outer core.

Pressure and Temperature Inside the Earth

In order to keep underground tunnels from collapsing under the pressure created by the weight of overlying rock, mining engineers must design sturdy support structures. It is no surprise that deeper tunnels require stronger supports: the downward push from the weight of overlying rock increases with depth, simply because the mass of the overlying rock layer increases with depth. In solid rock, the pressure at a depth of 1 km is about 300 atm. At the Earth’s centre, pressure probably reaches about 3,600,000 atm. Temperature also increases with depth in the Earth. Even on a cool winter’s day, miners who chisel away at gold veins exposed in tunnels 3.5 km below the surface swelter in temperatures of about 53°C (127°F). We refer to the rate of change in temperature with depth as the geothermal gradient. In the upper part of the crust, the geothermal gradient averages between 20°C and 30°C per km. At greater depths, the rate decreases to 10°C per km or less. Thus, 35 km below the surface of a continent, the temperature reaches 400°C to 700°C, and the mantle-core boundary is about 3,500°C. No one has ever directly measured the temperature at the Earth’s centre, but calculations suggest it may exceed 4,700°C, close to the Sun’s surface temperature of 5,500°C.

Geophysical and geological logging

Geophysical and geological logging

kacimo samy   5:45 ص

There is a wide range of instruments, geophysical logging tools, that are lowered down a borehole to record the physical and chemical properties of the rocks. These instruments are mounted on a device called a sonde that is lowered down the drill hole (on a wireline) once the drill string has been removed. Data from these instruments are recorded at the surface as the sonde passes up through the formations. An alternative technique is to fix a sonde mounted with logging instruments behind the drill bit and record data as drilling proceeds. The tools can be broadly divided into those that are concerned with the petrophysics of the formations, that is, the physical properties of the rocks and the fluids that they contain, and geological tools that provide sedimentological information. The interpretation of all the data is usually referred to as formation evaluation – the determination of the nature and properties of formations in the subsurface. Many of these tools are now used in combinations and provide an integrated output that indicates parameters such as sand:mud ratio, porosity, permeability and hydrocarbon saturation.

Petrophysical logging tools

Caliper log

The width of the borehole is initially determined by the size of the drill bit used, but it can vary depending on the nature of the lithology and the permeability of the formation. The borehole wall may cave in where there are less indurated lithologies such as mudrocks, and this can be seen as an anomalously wide interval of the hole. The caliper log can also detect parts of the borehole where the diameter is reduced by the accumulation of a mud cake on the inside: mud cakes are made up of the solid suspension in the drilling mud and form where there is a porous and permeable bed that allows the drilling fluid to penetrate, leaving the mud filtered out on the borehole wall.

Gamma-ray log

This records the natural gamma radioactivity in the rocks that comes from the decay of isotopes of potassium, uranium and thorium. The main use of this tool is to distinguish between mudrocks, which generally have a high potassium content and hence high natural radioactivity, and sandstone and limestone, both of which normally have a lower natural radioactivity. The gamma-ray log is often used to determine the ‘sand: shale ratio’ in a clastic succession (note that for petrophysical purposes, all mudrocks are called 'shales'). However, it should be noted that mica, feldspar, glauconite and some heavy minerals are also radioactive, and sandstones rich in any of these cannot always be distinguished from mudstones using this tool. Organic-rich rocks can also be detected with this tool because uranium is often naturally associated with organic matter. Mudrocks with high organic contents are sometimes referred to as 'hot shales' because of their high natural radioactivity. The spectral gamma-ray log records the radioactivity due to potassium, thorium and uranium separately, allowing the signal due to clay minerals to be separated from radioactivity associated with organic matter.

Resistivity logs

Resistivity logging tools are a range of instruments that are used to measure the electrical conductivity of the rocks and their pore fluids by passing an electrical current from one part of the sonde, through the rocks of the borehole wall measuring the current at another part of the sonde. Most minerals are poor conductors, with the exception of clay minerals that have charged ions in their structures. The resistivity measurements provide information about the composition of the pore fluids because hydrocarbons and fresh water are poor electrical conductors but saline groundwater is a good conductor of electricity. Resistivity logging tools are usually configured so that they are able to measure the resistivity at different distances into the formation away from the borehole wall. A microresistivity tool records the properties at the borehole wall, a ‘shallow’ log measures a short distance into the formation and a 'deep' log records the current that has passed through the formation well away from the borehole (these are sometimes called laterologs). Comparison of readings at different distances from the borehole wall can provide an indication of how far the drilling mud has penetrated into the formation and this gives a measure of the formation permeability. Induction logs are resistivity tools that indirectly generate and measure the electrical properties by the process of induction of a current.

Sonic log

The velocity of sound waves in the formation is determined by using a tool that comprises a pulsing sound source and receiver microphone that records how long it has taken for the sound to pass through the rock near the borehole. The sonic velocity is dependent upon two factors. First, lithologies composed of high-density material transmit sound faster than low-density rocks: for example, coal is a low-density material, basalt is high-density, and sandstones and limestones have intermediate densities. Second, if the rock is porous, the bulk density of the formation will be reduced, and hence the sonic velocity, so if the lithology is known, the porosity can be calculated, or vice versa. The velocities determined by this tool can be used for depth conversion of seismic reflection profiles.

Density logs

These tools operate by emitting gamma radiation and detecting the proportion of the radiation that returns to detectors on the tool. The amount of radiation returned is proportional to the electron density of the material bombarded and this is in turn proportional to the overall density of the formation. If the lithology is known, the porosity can be calculated as density decreases with increased porosity. The application of this tool is therefore very similar to that of the sonic logging tool.

Neutron logs

In this instance the tool has a source that emits neutrons and a detector that measures the energy of returning neutrons. Neutrons lose energy by colliding with a particle of similar mass, a hydrogen nucleus, so this logging tool effectively measures the hydrogen concentration of the formation. Hydrogen is mostly present in the pore spaces in the rock filled by formation fluids, oil or water (which have approximately the same hydrogen ion concentration) so the neutron log provides a measure of the porosity of the formation. However, clay minerals contain hydrogen ions as part of the mineral structure, so this tool does not provide a reliable indicator of the porosity in mudrocks or muddy sandstones or limestones.

Electromagnetic propagation log

The dielectric properties of the formation fluids are measured with this tool. It consists of microwave transmitters that propagate a pulse of electromagnetic energy through the formation and measures the attenuation of the wave with receivers. The measurements are related to the dielectric constant of the formation, which is in turn determined by the amount of water present. The tool therefore can be used to distinguish between oil and water in porous formations.

Nuclear magnetic resonance logs

Conventional porosity determination techniques do not provide information about the size of the pore spaces or how easily the fluid can be removed from those pores. Fluids that are bound to the surface of grains by capillary action cannot easily be removed and are therefore not producible fluids, and if pore spaces are small more fluid will be bound into the formation. The nuclear magnetic resonance (NMR) tool works by producing a strong magnetic field that polarises hydrogen nuclei in water and hydrocarbons. When the field is switched off the hydrogen nuclei relax to their previous state, but the rate at which they do so, the relaxation time, increases if they interact with grain surfaces. Measurement of the electromagnetic 'echo' produced during the relaxation period can thus be used as a measure of how much of the fluid is 'free' and how much of it is close to, and bound on to, grain surfaces. The tool operates by producing a pulsed magnetic field and measuring the echo many times a second.

Geological logging tools

Dipmeter log

The sonde for this tool has four or six separate devices for measuring the resistivity at the borehole wall. They are arranged around the sonde so that if there is a difference in the resistivity on different sides of the borehole, this will be detected. If the layering in the formations is inclined due to a tectonic tilt or crossstratification it is possible to detect the degree and direction of the tilt by comparing the readings of the different, horizontal resistivity devices. Hence this tool has the potential to measure the sedimentary or tectonic dip of layering.

Microimaging tools

These tools, often called borehole scanners, are also resistivity devices and use a large number of small receiving devices to provide an image of the resistivity of the whole borehole wall. If there are fine-scale contrasts in electrical properties, for instance where there are fine alternations of clay and sand, it is possible to image sedimentary structures as well as fractures in the rock. The images generated superficially resemble a photograph of the borehole wall, but is in fact a ‘map’ of variations in the resistivity.

Ultrasonic imaging logs

High-resolution measurements of the acoustic properties of the formations in the borehole walls are made by a rotating transmitter that emits an ultrasonic pulse and then records the reflected pulse with a receiver. The main use of this tool is to detect how uneven the borehole wall is, and this can be related to both lithology and the presence of fractures.

Borehole stratigraphy and sedimentology

Borehole stratigraphy and sedimentology

kacimo samy   5:35 ص

The interpretation of seismic reflection profiles provides a model for the stratigraphic and structural relationships that may exist in the subsurface. Data from these sources can provide some indicators of the lithologies in the subsurface, but a full geological picture can be obtained only by the addition of information on lithology and facies. This can be provided by drilling boreholes through the succession and either taking samples of the rocks and/or using geophysical tools to take detailed measurements of the rock properties. When a borehole is drilled there are a number of ways of collecting information from the subsurface, and these are briefly described below.

Borehole cuttings

In the course of drilling a deep borehole, a fluid is pumped down to the drill bit to lubricate it, remove the rock that has been cut (cuttings) and to counteract formation fluid pressures in the subsurface. Due to the weight of rocks above, fluids (water, oil and gas) trapped in porous and permeable strata will be under pressure, and without something to counteract that pressure they would rush to the surface up the borehole. The drilling fluid is therefore usually a ‘mud’, made up of a mixture of water or oil and powdered material, which gives the fluid a higher density: powdered barite (BaSO4) is often used because this mineral has a density of 4.48. The density of the drilling mud is varied to balance the pressure in the formations in the subsurface. The drilling mud is recirculated by being pumped down the inside of the drill string (pipe) and returning up the outside: because it is a dense, viscous fluid, it will bring the cuttings with it as it reaches the surface. The cuttings are filtered from the mud with a sieve and washed to provide a record of the strata that have been drilled. These cuttings are typically 1–5mm in diameter and are sieved out of the drilling mud at the surface. Recording the lithology of these drill chips (mud-logging) provides information about the rock types of the strata that have been penetrated by the borehole, but details such as sedimentary structures are not preserved. Microfossils such as foraminifera, nanofossils and palynomorphs can be recovered from cuttings and used in biostratigraphic analysis. There is usually a degree of mixing of material from different layers as the fluid returns up the borehole, so it is the depth at which a lithology or fossil first appears that is most significant.

Core

A drill bit can be designed such that it cuts an annulus of rock away leaving a cylinder in the centre, a core, that can be brought up to the surface. Where coring is being carried out the drilling is halted and the section of core is brought up to the surface in a sleeve inside the hollow drill string. As each section of core is brought to the surface it is placed in a box, which is labelled to show the depth interval it was recovered from. Recovery is often incomplete, with only part of the succession drilled preserved, and the core may be broken up during drilling. The core is then usually cut vertically to provide a smooth-surfaced slab of rock that is typically 90 mm to 150 mm across, depending on the width of the borehole being drilled. Cores cut in this way provide a considerable amount of detail of the lithologies present, the small-scale sedimentary structures, body and trace fossils. In exploration for oil and gas and in the development of fields for hydrocarbon production, cores are cut through ‘target horizons’, that is, parts of the succession that have been identified from the interpretation of seismic interpretation as likely source rocks, or, more importantly, reservoir bodies. Core is usually only cut and recovered through these parts of the stratigraphy: the rest of the succession has to be interpreted on the basis of geophysical wireline logs. However, continuous cores may be cut through successions that cannot be interpreted satisfactorily using geophysical information alone, as can occur when the properties of the rock units do not allow differentiation between different lithologies using wireline logging tools. In contrast to oil and gas exploration, coal and mineral exploration normally involves taking a complete core through the section drilled. The width of the core that is cut is smaller, often just 40mm, and the core is not split vertically. The small size and the curved surface of the core may make it more difficult to recognise sedimentary structures than in the conventional, larger, split core used in oil and gas exploration, but the continuous core provides good vertical coverage of the drilled succession.

Core logging

The procedure for recording the details of the sedimentary rocks in a core is very similar to making a graphic sedimentary log of a succession exposed in the field. Core logging sheets are similar in format to field logging sheets, and the same types of information are recorded (lithology, bed thickness, bed boundaries, sedimentary structures, biogenic structures, and so on). The scale is usually 1:20 or 1:50. In some ways recording information about strata from core is easier than field description. If the core recovery is good then there will be an almost complete record of the succession, including the finer grained lithologies. Weathering of mudrocks in the field usually means that they are less well preserved than the coarser beds, but in core this tends to be less of a problem, although weaker, finer grained beds will often break up more during the drilling. The main limitations are those imposed by the width of the core. It is not possible to see the lateral geometry of the beds and recognise features such as channels easily, and only parts of larger scale sedimentary structures are preserved. On the other hand, the details of ripple-scale features may be more easily seen on the smooth, cut surface of a core. Palaeocurrent data can be recorded from sedimentary structures only if the orientation of the core has been recorded during the drilling process, and this is not always possible. The other, not insignificant, difference between core and outcrop is that the geologist can carry out the recording of data in the relative comfort of a core store, although it is unlikely to be such an interesting environment to work in as a field location in an exotic place. Not all cores pass through the strata at right angles to the bedding. If the strata are tilted then a vertical drill core will cut through the beds at an angle, so all bed boundaries and sedimentary structures observed in the core will be inclined. During the development phase of oil and gas extraction, drilling is often directed along pathways (directional drilling) that can be at any angle, including horizontal. Interpretation of inclined and near-horizontal cores therefore requires information about the angle of the well.

Marine and land seismic aquisition

Marine and land seismic aquisition

kacimo samy   10:38 ص

Marine aquisition

Marine seismic acquisition is generally accomplished using large ships with one or multiple air-gun arrays for sources. Air-guns are deployed behind the seismic vessel and generate a seismic signal by forcing highly pressurised air into the water at a given interval. Receivers are towed behind the ship in one or several long streamer(s) that are several kilometres long. Marine receivers are composed of piezoelectric hydrophones, which respond to changes in water pressure. In marine acquisition, seismic vessels sail along predetermined patterns of parallel circuits. The length of straight segments is calculated from fold plots, and must include additional length run in and run out to allow the cable to straighten after each turn. Marine seismic acquisition is faster than conventional land seismic because it does not require jug hustlers to lay and pick up geophones. The advance to 3D seismic acquisition and imaging of the subsurface, introduced in the 1980s, was perhaps the most important step in seismic exploration. The 3D seismic images began to resolve the detailed subsurface structural and stratigraphic conditions that were missing or not discernable from 2D seismic data. With 3D seismic acquisition potential reservoirs are imaged in three dimensions, which allows seismic interpreters to view the data in cross-sections along 360◦ of azimuth, in depth slices parallel to the ground surface, and along planes that cut arbitrarily through the data volume. Information such as faulting and fracturing, bedding plane direction, the presence of pore fluids, complex geological structure, and detailed stratigraphy are now commonly interpreted from 3D seismic data sets. In 2D marine data acquisition a single streamer is deployed, whereas in 3D acquisition multiple streamers are towed behind the boat. 3D acquired data can be processed in a more consistent manner and be further manipulated using modern visualisation tools. In 2D data acquisition the data collection occurs along a line of receivers. The resultant image represents only a section below the line. Unfortunately, this method does not always produce a clear subsurface image. 2D data can often be distorted with diffractions and events produced from offline geologic structures, making accurate interpretation difficult. Because seismic waves travel along expanding wavefronts they have a surface area. A truly representative image of the subsurface is only obtained when the entire wavefield is sampled. A 3D seismic survey where the vessel is towing many streamers (up to 20) at the same time with multiple arrays of airguns is more capable of accurately imaging reflected waves because it utilises multiple points of observations. In the case of 3D survey we have seen that seismic data are sampled from a range of different angles (azimuth) and source-receiver distances (offsets). After seismic processing the data can then be represented as 3D volume images of the subsurface. In 2D processing, traces are collected into CMP gathers, while in 3D, traces are collected into common-cell gathers (binning). To perform 3D binning, a grid is first superimposed on the survey area. This grid consists of cells with dimensions of half the receiver group spacing in the in line direction, equivalent to the CMP spacing in 2D processing, and the line spacing in the cross line direction. In reality, midpoint distributions within a cell are not necessarily uniform since cable shape varies from shot to shot and line to line. Such side drift of the cables is called feathering. Another advanced technique of marine seismic is ocean bottom seismic (OBS) acquisition. It gives the possibility of direct measurement of S-wave data in addition to P-wave data by using ocean bottom cables that have three component geophones (3C) and a hydrophone in addition (thus 4C in total). The 4C cable can be up to 6 km long with 240 stations (i.e. 960 channels since 4C). A typical OBS layout involves 4 or more cables. The optimal choice of acquisition geometry for a 4C survey hinges on both geophysical and financial considerations. Most designs can be classified as either patch or swath. In swath designs, the source lines are parallel to receiver lines, while in patch designs, source lines are perpendicular to receiver lines. Patch design produces seismic data of a relatively wide range of azimuths, whereas the swath design produces data of a limited or narrow range. Moreover, the swath design offers a more uniform sampling of offsets with better near-offset coverage. The use of 4C OBS recording has several advantages over conventional towed streamer technology, which includes:

1. Dual-sensor summation (3C geophone + hydrophone) for the suppression of receiver-side multiples. 
2. Utilising P–S wave conversions for enhanced imaging. 
3. Attenuation of free surface multiples when combined with towed streamer recording.

A comparison of the migrated P–S stack versus the P–P stack. The P–S stack is produced from OBS converted wave data whereas the P–P stack is produced from 3D towed streamer P-wave data. From this comparison it is clear that OBS data can be used to successfully image through a gas chimney.

Land aquisition

A complication in land acquisition is that, unlike marine data, a seismic line is rarely shot in a straight line because of the presence of natural and man made obstructions such as lakes, buildings and roads. The shot points and the receivers may be arranged in many ways. Many groups of geophones are commonly used on a line with shot points at the end or in the middle of the receiver array. The shot points are gradually moved along a line of geophones. The variations in ground elevation in land acquisition causes sound waves to reach the recording geophones with different travel time. The Earth’s nearsurface layer may also vary greatly in composition, from soft alluvial sediments to hard rocks. This means that the velocity of sound waves transmitted through this surface layer may be highly variable. Static corrections a bulk time shift applied to a seismic trace are typically used in seismic processing to compensate for these differences in elevations of sources and receivers and near-surface velocity variations.

Sources and receivers in seismic aquisition

Sources and receivers in seismic aquisition

kacimo samy   8:24 ص

It is essential at this point to understand the principle of seismology. Seismology is based on the transmission of sound waves by the rocks of the crust. Strong earthquakes create pressure waves (natural sources of seismic waves) that are transmitted through the entire Earth and detected by seismographs (receivers) on the other side. Seismic exploration, however, as employed by the petroleum geologist, makes use of artificially generated pulses (seismic sources). The principle is simple; an impulse source sends acoustic energy into the Earth. This energy propagates in many directions and is reflected and refracted when it encounters boundaries between two layers. Sensors (seismic receivers) placed on the surface measure the reflected or refracted acoustic energy. These artificial sources are much weaker than the natural seismic source (earthquake) but are more focused towards areas of specific stratigraphic interest.

Seismic Sources

Different seismic sources are usually used in land and marine acquisitions. In marine environments seismic energy is normally generated using arrays of air-guns, whereas in land seismic one often uses explosives or vibrators. An air-gun is a device that releases highly compressed air (at typically 2,000–5,000 psi) into the water surrounding the gun. A vibrator is an adjustable mechanical source that delivers vibratory seismic energy into the ground. A vibrator source sends a controlled-frequency sweep into the ground. The recorded data are then convolved with the original sweep to produce a usable signal. Dynamite a combination of explosive and detonator, is used as a seismic source. The detonator helps to ignite the explosives. When dynamite ignites, a shock wave propagates with a speed of 3,000–10,000 m/s. It provides an impulsive energy that can be converted into ground motion. It is customary to drill a hole to load dynamite and fill it with heavy mud before shooting. Dynamite can generate usable signal strengths and a bandwidth that covers a wide spectrum of seismic energy. It includes a variety of energy sources based on varying explosive out put parameters to meet geological and climatic conditions.

Seismic Receivers

Hydrophones and geophones serve as receivers for seismic signals. The hydrophone is a device designed for use in detecting seismic energy in the form of pressure changes in water during marine seismic acquisition. It measures pressure variations with the aid of piezoelectric material, which generates a voltage upon deformation. The two piezoelectric elements in one hydrophone are connected and polarised so that voltages due to pressure waves (returning signal) add and voltages due to one-directional acceleration will cancel. In this way the influence of movements due to currents, wave action and so on will be minimised. Hydrophones are combined to form streamers that are towed by seismic vessels or deployed in a borehole. A typical length of a streamer is about 4-6 km where a single receiver section is typically 75 m long and contains 96 hydrophones which are grouped in arrays of a pre-defined length, mostly 12.5 or 25 m. The geophone is a device used in surface seismic acquisition, both onshore and on the seabed offshore, that detects ground velocity produced by seismic waves and transforms the motion into electrical impulses. Geophones, unlike hydrophones, detect motion rather than pressure. Conventional seismic surveys on land use one geophone or a group of geophones per receiver location to detect motion in the vertical direction. The three-component (3C) geophone is used for direct measurements of shear waves at the seafloor. An essential feature of a seafloor seismic acquisition is the four-component (4C) detector unit, which includes a hydrophone and a three-component (3C) geophone. The hydrophone and the vertical geophone measure pressure waves, whereas the two extra horizontal geophones measure particle velocity associated with shear wave energy.

Seismic aquisition

The seismic survey is an essential part of the whole cycle of petroleum exploration and production. Seismic surveys are carried out on land and in transition zone, shallow marine and marine environments in different ways. The basic principle is an impulse source such as dynamite, air-gun or vibrator that sends acoustic energy into the Earth. This energy propagates in many directions. Downward travelling energy reflects and refracts when it encounters boundaries between two layers with different acoustic properties. Sensors or geophones placed on the surface measure the reflected acoustic energy, converting it into an electrical signal that is displayed as a seismic trace. The typical recorded seismic frequencies are in the range of 5–100 Hz. P-waves are the waves generally studied in conventional seismic data. P-waves incident on an interface at other than normal incidence angle can produce reflected and transmitted S-waves. S-waves travel through the Earth at about half the speed of P-waves and respond differently to fluid-filled rocks, and so can provide different additional information about lithology and fluid content of hydrocarbon-bearing reservoirs. The recorded seismic trace is a convolution (∗) of the source signal and the reflectivity sequence of the Earth plus noise. A seismic trace can simply be expressed by the equation below where multiples are not considered. Transmission losses and geometric spreading are not included and the frequency dependent absorptions are also ignored in the equation.

S = W ∗ R + Noises 

where S is the recorded seismic trace, R is the reflectivity and W is the wavelet. A wavelet is a kind of mathematical function used to divide a given function into different frequency components and study each component with a resolution that matches its scale. Accurate wavelet estimation is absolutely critical to the success of any seismic inversion. The inferred shape of the seismic wavelet may strongly influence the seismic inversion results and therefore subsequent assessments of the reservoir quality. Attenuation (amplitude loss) of seismic waves is an important phenomenon and caused by three major factors: 

• Geometric spreading: progressive diminution of amplitude (proportional to the inverse of propagation distance) caused by the increase in wavefront area. 
• Intrinsic attenuation: energy losses due to internal friction. 
• Transmission losses: reduction in wave amplitude due to reflection at interfaces.