Carbonate Petrography

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

25 strangest Geologic Formations on Earth

The strangest formations on Earth.

What causes Earthquake?

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

The Geologic Column

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

Folds and Foliations

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

How are sedimentary basins formed?

Sedimentary basin: A low area in the Earth’s crust, of tectonic origin, in which sediments accumulate. Sedimentary basins range in size from as small as hundreds of meters to large parts of ocean basins. The essential element of the concept is tectonic creation of relief, to provide both a source of sediment and a relatively low place for the deposition of that sediment. Tectonics is the most important control on sedimentation; climate is a rather distant second. The important effects of tectonics on sedimentation, direct or indirect, include the following:
  •  nature of sediment
  • rate of sediment supply 
  • rate of deposition 
  • depositional environment 
  • nature of source rocks 
  • nature of vertical succession

The only basins that are preserved in their entirety are those that lie entirely in the subsurface! Basins exposed at the surface are undergoing destruction and loss of record by erosion. So there’s an ironic trade-off between having more complete preservation in the subsurface but less satisfactory observations. How do you gather data on sedimentary basins?.

Master cross sections: With the present land surface as the most natural datum, construct several detailed physical cross sections through the basin to show its geometry and sediment fill.

Stratigraphic sections: Construct a graph, with time along the vertical axis, showing the time correlations of all the major rock units along some generalized traverse across the basin. Such a section includes hiatuses, during which there was non deposition or erosion.

Isopach maps: With some distinctive stratigraphic horizon near the top of the section as datum, draw a contour map showing isopachs (isopachs are loci of equal total sediment thickness) in the basin.

Lithofacies maps: For one or a series of times, draw a map showing distribution of sediment types being deposited at that time.

Ratio maps: Compute things like sand/shale ratio, integrated over the entire section or restricted to some time interval, and plot a contour map of the values.

Paleocurrent maps: For one or a series of times, draw a map showing the direction of paleo-currents in the basin at that time.

Grain-size maps: For the entire basin fill, averaged vertically, or for some stratigraphic interval or time interval, draw a map that shows the areal distribution of sediment grain size. This is especially useful for conglomeratic basins.

In one sense, the origin of sedimentary basins boils down to the question of how relief on the Earth is created. Basically, there are only a few ways, described in the following sections.


On a small scale, hundreds to thousands of meters laterally, fault movements can create relief of hundreds to thousands of meters, resulting in small but often deep basins (some of these are called intermontane basins). Along strike-slip faults. can produce small pull-apart basins; more on them later. Relief of this kind is on such a small scale that it tends not to be isostatically compensated. It’s like setting a block of granite out on your driveway; the flexural rigidity of your driveway is great enough compared with the imposed load that the granite block is prevented from finding its buoyant equilibrium position.


Basin relief can be created mechanically on a regional scale in two very important ways: thermally or flexurally, or by a combination of those two effects). Each of these is discussed briefly below. Keep in mind that basins can also be made just by making mountain ranges, on land or in the ocean, by volcanism. 


If the lithosphere is heated from below, it expands slightly and thus becomes less dense. This less dense lithosphere adjusts isostatically to float higher in the asthenosphere, producing what we see at the Earth’s surface as crustal uplift. If the lithosphere cools back to its original temperature, there’s isostatic subsidence back to the original level.

 But suppose that some erosion took place while the crust was elevated. The crust is thinned where the erosion took place (and thickened somewhere else, where there was deposition; that might be far away, at the mouth of some long river system), so when the crust cools again it subsides to a position lower than where it started, thus creating a basin available for filling by sediments.

 But the magnitude of crustal lowering by this mechanism is less than is often observed in basins thought to be created thermally. It has therefore been proposed, and widely accepted, that in many cases extensional thinning of the lithosphere accompanies the heating. Then, upon re-cooling, the elevation of the top of the lithosphere is less than before the heating and extension. This kind of subsidence has been invoked to explain many sedimentary basins.        



Another important way to make basins is to park a large load on some area of the lithosphere. The new load causes that lithosphere to subside by isostatic adjustment. But because the lithosphere has considerable flexural rigidity, adjacent lithosphere is bowed down also. The region between the high-standing load and the lithosphere in the far field (in the parlance of geophysics, that just means far away!) is thus depressed to form a basin. This model has been very successful in accounting for the features of foreland basins, which are formed ahead of large thrust sheets that move out from orogenic areas onto previously undeformed cratonal lithosphere.

How to construct a cross section using the Kink method?

Kink method: The basic method is to allow each dip measurement to define a zone where the dip is constant. The boundaries of the dip zones are the lines that bisect angles between adjacent dips. The example below begins with three different ways to find the bisector.
We first find the line L12 that bisects the angle between dips 1 and 2. 
One way is to construct lines perpendicular to dips 1 and 2 (in red). They intersect at C12. Bisect the angle formed by the two perpendiculars (shown in green).

Extend the dips out to L12. When you cross the line, continue with the dip on the opposite side. That is, use dip angle 1 on the left side of line L12 and dip angle 2 on the right.

Here's another way: halfway between datum points 2 and 3, draw lines parallel to dips 2 and 3 (shown in red). Bisect the angle to find line L23.

Extend dips 2 and 3 to L23, then continue on the opposite side of the line with the appropriate dip. For example, extend dip 2, then where it intersects L23, continue with dip 3.
We can also extend the line that originates at datum point 1. Between L12 and L23 it has dip 2, then on the right of L23 it has dip 3.

Yet another approach is to average two adjacent dips. The desired line passes through the midway point between the two datum points. Draw a vertical line through that point and measure the average angle from the vertical line. Pay careful attention to dip direction so that the line has the proper slope.

Lines left of L34 have dip 3, to the right they have dip 4. Continue the lines that originated at datum points 1 and 2 by extending them through L34 with the appropriate dips.

L45 constructed. Lines are extended to L34 and the dip beyond is indicated. Dips 4 and 5 are very close to already existing lines and are unlikely to give any new information, so we don't need to construct lines through dips 4 and 5 for now.

L56 constructed. The rocks at datum point 6 are stratigraphically lower than any we have so far seen, so we extend dip 6 to line L56, then continue to the left parallel to the previously drawn lines.

L67 constructed. The rocks at datum point 7 are even lower than those at datum point 6, so we also extend dip 7 to line L67, then continue to the left parallel to the previously drawn lines.

L78 constructed. The rocks at datum point 8 are still lower , so we also extend dip 8 to line L78, then continue to the left parallel to the previously drawn lines.

L89 constructed. Datum point 9 lies between two already constructed lines and is not likely to tell us anything we don't already know. For now we omit drawing a line through it.

L910 completed. Since point 10 falls close to an already drawn line, there is no real need to construct another line for it, either.

Here we have indicated the stratigraphy. It is virtually certain  when you draw a cross section using strictly geometric methods that the contacts will not match exactly with their predicted positions.
What we need to do now is redraw the folds so the cross-section matches both the dips and the stratigraphy.

Here the cross section lines are subdued. 
Most of the time you can modify the fold shapes by hand to match the stratigraphy without too much trouble. Modified contacts are in black. 

How to construct a geological map?

Before constructing a geological map it is important to be clear out about the aims of the investigations, for this decision will guide the choice of map scale and control the nature of the techniques which are needed to cover the area in detail to resolve the problem.

Well what is a geological map?.

A geologic map portrays the distribution of rocks, deposits, or other geologic features in a specified area. Each consolidated rock type that can be distinguished by similar characteristics is categorized into a mappable unit, or formation. Unconsolidated deposits such as landslides and stream alluvium also are designated on our geologic maps. Unique colours, patterns, and labels are used to differentiate each unit on the map. Colours are chosen by the age of the rocks being described. For example, rocks from the Jurassic Period are coloured in shades of green, and Quaternary deposits are coloured shades of yellow. Labels designate the age and name of the formation or deposit. A geologic map is typically printed over a topographic base map.

Many different types of lines and symbols are found on geologic maps. The most prevalent are thin black lines that depict the contacts between two different mappable units. Line width and colour are used to differentiate other types of features such as faults and folds. Geologists collect structural information describing whether the rock layers are tilted or not. Most sedimentary and volcanic rocks were originally deposited horizontally. Therefore, the tilt (dip) of layered rock units may provide key information to understanding whether non-horizontal units were deformed by faulting or folding. This type of structural information also helps geologists build interpretive cross-sections that depict how the map units may look in the subsurface. Geologic maps provide a wealth of information to all who use them. They are important for geologic hazard detection and mitigation, mineral and groundwater resource evaluation, and provide enjoyment for the casual roadside geologist. Geologic maps help us understand the earth on which we walk, and give us a greater appreciation for the geology around us.

First of all you have to mark your area in Google Earth and save it as KMZ file by right clicking the area folder and then click save places as KMZ.

Location of study area.

Then Mark your formations with polygon and save it as KMZ file by right clicking each formation.

Formations polygons being made.
Now open each file in global mapper and export it separately into Arc GIS shape file by clicking File > Export Vector format > Shape file.

Location area in global mapper.
First polygon in global mapper.

Second polygon in global mapper.

Third polygon in global mapper.

Now you can open it in Arc GIS as shape file for GIS already been created. Now it will look like this in GIS.

All data when loaded in Arc GIS.

So now what you have to do  when all being loaded?.

Simply complete your map, You can add the formations name which will self name at what you saved in the Google Earth. You can label by simply clicking each polygon by right click and label features. Now what else is missing?. This map is already being geo-referenced from the Google Earth moreover click on the insert and add legends, North arrow, Co-ordinates etc what ever is your need suit yourself. This is how you can make geological map using Arc GIS. You can also use Coral Draw for constructing map.

What are earthquakes and what causes them?


An earthquake is the brittle, sudden failure of the earth's crust or mantle. Earthquakes are caused by several factors however the common element is that stress builds in rocks until the yield strength of the rock is exceeded, at which point rupture occurs. The relative movement between the major tectonic plates is responsible for the stress build-up that causes the vast majority of earthquakes.

To understand the earthquake process some simple physical quantities should be defined first.

Stress is the force per unit area.
Elastic deformation: A material changes shape when stressed but after the stress is removed, it returns to its original shape. The bonds between the molecules and atoms of an elastically behaving material, when stressed, stretch and bend, but retain memory of their original configuration. Once the stress is released, the stored energy is released and the material returns to its original shape examples are rubber band or super ball and rocks
Brittle deformation: Rupture occurs in response when stress that is exerted on the material exceeds that materials yield strength example are glass, ceramic and rocks
Plastic deformation: Flow occurs in response to stress and material does not return to original shape after applied stress is removed. Just like a Play Dough which can be shaped as liked. Some rocks (rock salt or halite) and other evaporites flow when subjected to stress.

How does elastic rocks break?. 

It can be best understood by the following example of a plywood. When the strip of plywood is subjected to a bending force, it deforms elastically at first. If the bending force is released, the wood returns to its original unbent shape. However, if the bending exceeds the yield strength of the weakest part of the wood, that part will rupture. Once the rupture occurs, we all hear the cracking noise - this represents the propagation of acoustic energy through the air due to the rupture. In other words, the rupture releases energy into the surrounding medium and this energy spreads away from the point of rupture. The farther that one is located from the point of the rupture, the softer the rupture noise is because that finite amount of energy released by the cracking wood is being spread over a larger volume as it moves away from the source.

Fault breaks which results in earthquake

So how does earthquake occurs in the crust?

In the earth's crust and in particular, within the fault zones that accommodate the motion between the rigid interiors of plates, the crust deforms elastically between earthquakes. The faults have geometric irregularities (bends) that prevent the crust on either side of the fault from slipping smoothly (creeping) in response to the steady state motion of the plates on either side of the fault. Because friction prevents steady state slip along a fault, rocks near the fault deform elastically in response to plate motion far from the fault. Once the amount of elastically stored energy exceeds the strength of the weakest area of rocks along a fault, that patch of the fault ruptures. At the point of rupture, rocks on either side of the fault slip to their new location and in the process, release lots of stored energy that propagates away from the point of rupture. A small rupture in one area of a fault can place a sudden strain on a nearby, more strongly locked section of the fault and cause that section of the fault to rupture, too. Thus, one earthquake can trigger another. Faults often have bends; the rocks on the fault face can have different frictional and elastic properties; fluids may lubricate the fault; and other nearby faults may change the local stresses.

Once an earthquake has occurred along a section of a fault, much of the stress on those rocks is relieved. However, since steady state plate motion is still occurring, stress immediately begins to build again, leading to the earthquake cycle in which repeat earthquakes occur along sections of a fault. The frequency and strength of earthquakes along a given fault depends on how quickly the stress builds, how weakly or strongly the fault is locked in a particular region, and interactions with other nearby faults that are also responding to the stress build up. This makes it difficult to to model the earthquake cycle.

Rupture and propagation of seismic energy

To understand the propagation of seismic waves one can demonstrate it by throwing a stone into a pond. Well ripples on the pond carries energy away from the point of impact. Some of the energy is also carried down into the pond as sound, which we could hear if we were beneath water when the stone was thrown. In a similar fashion, during an earthquake rupture, two broad categories of seismic waves are generated.
  • Body waves, which carry seismic energy through the interior of the earth
  • Surface waves, which carry seismic energy along the surface.
  • Body waves can be further sub-divided as follows:
  • P wave (primary) is Compressional. Particles displaced in direction of energy propagation
  • S wave (secondary) is Shearing. Particles displaced perpendicular to direction of energy propagation.
Surface waves, which cause the earth's surface to roll as they pass by are often responsible for the majority of earthquake damage. Surface wave amplitudes can reach several meters, meaning that during a large earthquake, one end of your house could be in the trough of a surface wave several meters beneath the other end of your house which could be surfing on the crest of a surface wave. Surface waves travel slowly often take several minutes or longer to travel tens of miles. Body waves arrive within seconds but aren't as likely to cause major shaking.

Why are seismic waves useful?

Seismic waves are useful for locating earthquakes, determining the amount of energy that was released, and determining what type of fault slip occurred. Seismologists routinely exploit this information using a global network of seismographs that continuously feed their readings into several analysis centres. Earthquake locations (epicentres) and magnitudes are typically available less than an hour after an earthquake. 

To find the location, three things required to completely describe the location of an earthquake 
  • Its latitude 
  • Its longitude 
  • Its depth. 
These three together describe the earthquake focal point, which is the point within the earth where an earthquake started to rupture a fault. The point on the earth's surface directly above the focus is called the epicentre.
Magnitude: The magnitude of an earthquake measures the total amount of ground shaking produced near the epicentre. There are many scales to measure the magnitude of an earthquake but the most used one is Richter scale Richter magnitudes vary from 1 to about 9, with 1 being very small and 9 being enormous. In general, an increase of 1 point in the magnitude represents a 10X increase in the amount of ground motion and a 31X increase in the amount of energy release.
Intensity: An alternative way to measure the size of an earthquake is by its effect on humans and surface features such as buildings. This technique has shortcomings because it depends on the often subjective observations of individuals. However, for earthquakes that occurred before regular instrumental recording made it possibly to routinely estimate earthquake magnitudes, estimates of intensity are the only way to locate epicentres and determine how large the earthquake was. 

Earthquake Risk factors

  • Fault movement: direct breakage of structure built on fault trace.
  • Ground Shaking: ground vibration caused by seismic waves travelling away from focus.
  • Landslide: ground shaking can induce failure of weak slopes.
  • Liquefaction: ground shaking of wet soil can induce creep of soil.
  • Tsunami: disturbance of sea floor causing seismic sea wave.
  • Fire: rupture of gas lines etc.

How do mountains build?

Plate moves around the Earth's surface which exerts powerful lateral forces on rocks. The response of the crust to those forces gives rise to deformation on a large scale, particularly along plate boundaries. For example, dozens or hundreds of large-scale faults can form in zones of plate convergence, resulting in a broad and high mountain belt. The geologic processes that can result in mountain building.
  • Crustal shortening/thickening in response to convergence of a subducting plate.
  • Continental collision.
  • Uplift of sediments accreted by subduction.
  • Volcanism.
  • "Corner" accretion/uplift due to along-trench rafting of terrain.
  • Heating or cooling of lithosphere by the underlying mantle (and hence uplift or subsidence).
  • Crustal extension.
Geometrically, if two plates are colliding, they can respond as follows:
  • One plate can subduct into the mantle.
  • One or both plates can undergo shortening and hence uplift and crustal thickening.
  • One or both plates can undergo lateral extrusion (i.e. part of the plate escapes the collision zone by extruding sideways).
Let's consider the first two cases.
Suppose an oceanic plate converges with a continental plate. The oceanic plate, being more dense, subducts into the mantle. If all convergence (100%) is accommodated by subduction, the overlying continent will remain undeformed and should undergo no net uplift (although topography will undoubtedly develop as arc volcanoes appear). Suppose that not all convergence is accommodated by subduction and a few percent of the convergence instead causes the upper plate to shorten. How much uplift is implied?

For 100 millimetres per year of convergence (a typical plate velocity), if 5% is accommodated by long-term shortening of the overlying plate, then the overlying plate will shorten at a rate of roughly 5 millimetres per year.

Geometrically, the 5 mm/yr of horizontal shortening translates into vertical motion as shown above, with the only dependence being on the dip of the fault along which the continent is shortening. For low angle faults (10 degrees), uplift will occur at rates of about 1 mm/yr (1 kilometre per million years). Over 5 million years, this amounts to uplift of 5 kilometres or 3 miles (15,900 feet). Over the same period, the two plates will have converged some 5 million years x 100 kilometres per million years or 500 km. The total uplift then is only a small percentage, 1%, of the total horizontal motion.

Continent-continent collisions are not usually accompanied by subduction because both plates are too buoyant to be thrust deeply into the mantle. The amount of crustal thickening and uplift in such a collision can thus be much greater than for an ocean-continent collision. For example, over the past 40 million years, continental India has driven northward into continental Eurasia across the Himalayas mountain belt at a rate of about 40-50 millimetres per year. If you work it out, this implies that the two plates have somehow shortened by more than 1500 kilometres (about 900 miles) across the Himalayas. Since subduction is not occurring, the shortening has been accommodated by mountain building and lateral escape.

Factors that complicate calculations of total uplift

Plate convergence of hundreds or even a thousand kilometres or more over millions of years might be accompanied by only several kilometres of uplift. This uplift, while spectacular to the eye, is merely a small part of the displacement "budget", which is largely dominated by horizontal motion.

Many important aspects of mountain building are ignored and thus cannot predict total topography given total horizontal motion.

Why is this the case?. It is because of the following:

  • Fault dips are rarely well known and faults are often curved, with their dips increasing toward the surface.
  • Erosion/mass wasting removes material from the upper reaches of uplifting regions, sometimes nearly as quickly as the region is going up!
  • Vertical uplift must fight the downward pull of gravity. In extreme cases, no amount of horizontal convergence is capable of causing further uplift.
  • Crustal uplift often coincides with subsidence through a process called isostatic adjustment. This is analogous to climbing into a boat, which sinks lower into the water once your weight is added to that of the boat. Your net height above the water surface is equal to your height on land minus the amount the boat sinks. Continents similarly sink into the underlying mantle when substantial topographic loads are added to their surface.
  • Lateral escape of crust sometimes accommodates a significant fraction of plate convergence, leaving less available for inducing vertical uplift.
  • Uplift can occur simultaneously or sequentially along many faults that exist in a broad zone of deformation between two converging plates. Relating the total horizontal convergence to the total uplift caused by the convergence then requires measurements across many faults.