SAARC Workshop on

Geophysical Techniques for

Exploration of

Natural Resources

18-22, October, 2010

Khalid Amin Khan aminkhan@k-tron.net

Oil & Gas Training Institute, OGDCL

Islamabad, Pakistan

This manual is a subset of my original training manual

Seismic Methods

K. A Khan, 2009

It is a supplement to

Seismic Methods, Digital Courseware Series, 2nd Edition

K. A Khan, 2009

“I do not know what I may appear to the world, but to

myself I seem to have been only a boy playing on the sea-

shore, and diverting myself in now and then finding a

smoother pebble or a prettier shell than ordinary, whilst

the great ocean of truth lay all undiscovered before me.”

Isaac Newton

Training of Professionals from SAARC Countries

Geophysical Techniques for Exploration of Natural Resources By

Khalid Amin Khan, Dy.Chief Geophysicist

Oil & Gas Training Institute, OGDCL, Islamabad

Schedule

Day-1

• Physics and Electromagnetic Spectrum

• Imaging the Invisible

• Overview of Geophysical Methods

• Electrical Resistivity Methods

� Resistivity Meter

� Schlumberger Configuration

� Wenner Configuration

� Vertical Electrical Sounding Curves

Day-2

• Gravity & Magnetic Methods

� Gravity Meter & Magnetometer

� Gravity Field Correction: Free Air and Bouguer Anomaly

� Terrain Corrections

� Regional and Residual Separation

� Gravity Modelling using Talwani Method

� Magnetic Data Processing

� 2D Grid Processing

• Seismic Waves and Rock Physics

� Types of Seismic Waves

� Seismic Velocities

� Engineering Properties and Rock Physics

Day-3

• Seismic Refraction Methods

� First Breaks: Direct and Refracted Waves

� Automated First Break Picking

� TX-Graphs and Layer Velocity and Depth Computation

� Low Velocity Weathered Layers and Statics Computation

� Interpolation and Gridding

� Uphole Logging Methods

• Seismic Reflection Data Acquisition

� Geophones as Transducers

� Seismic Recorder

� Multiplexing and De-multiplexing

• Seismic Noise

� Coherent and Incoherent Noise

� Geophone Arrays

� Low and High Cut Filters

� Stacking to remove Incoherent Noise

Day-4

• Seismic Data Processing I

� Data Processing System Environment

� Processing Tasks and Job Control Language

� Basic Processing Flow

� Gains, Spherical Divergence

� Band Pass Filter

� Deconvolution

• Seismic Data Processing II

� Dynamic Corrections / Normal Moveout

� Velocity Analysis / Constant Velocity Stack

� Stacking: Raw, Brute Stack, Residual Statics & Migration

Day-5

• Seismic Resolution

� Temporal Resolution: Frequency and Bandwidth

� Spatial Resolution: Picket Interval and Fresnel Zone

� Phase Uncertainty

� Signal to Noise Ratio

• Seismic Interpretation

� Components of a Base Map

� Seismic Section: Display Modes, Vertical and Horizontal Scales

� Components of a Petroleum System

� Marking Horizons and Faults

� Auto Tracking Horizons

� Posting Data to Base Map

� Contouring

� 3D Seismic Cube: Inline Section, Cross-line Section & Time Slice

� Sonic and Bulk Density Logs for Synthetic Seismogram

� Seismic Modeling

� Seismic Velocities and Time to Depth Conversion

Contents

Module 1: Physics and Electromagnetic Spectrum

1.1 Basic Foundations of Physics

1.2 Imaging Principle

1.3 Imaging the Invisible

1.4 Fundamental Laws of Wave Propagation

Module 2: Electrical Resistivity Methods

2.1 Electrical Resistivity Methods

2.2 Resistivity Meter

2.3 Electrical Resistivity Surveying

2.4 Electrode Geometries

2.5 Resistivity Interpretation

Module 3: Gravity & Magnetic Methods

3.1 Gravity and Magnetic Prospecting

3.2 Gravity Method

3.3 Gravimeters

3.4 Gravity Surveying and Corrections

3.5 Regional Residual Separations

3.6 Gravity Modeling

3.7 Magnetic Method

3.8 Magnetometers

3.9 Magnetic Surveying and Corrections

3.10 2D Grid Processing

Module 4: Seismic Waves and Rock Physics

4.1 Seismic Waves

4.2 Types of Seismic Waves

4.3 Uses of Seismic Waves

4.4 Stress and Strain

4.5 Elasticity & Stiffness

4.6 Hooks Law of Elasticity

4.7 Elastic Moduli

4.8 Computing Density/Moduli from Seismic Velocities

Module 5: Seismic Refraction Methods

5.1 Snell’s Law

5.2 Seismic Refraction Method

5.3 Seismic Refraction Data Acquisition

5.4 First Breaks

5.5 Time-Distance Graphs

5.6 Statics Corrections

5.7 Limitations of Seismic Refraction Method

Module 6: Seismic Reflection Data Acquisition

6.1 Digital Sampling and Aliasing

6.2 Seismic Recorder

6.3 Seismic Sources

6.4 Fold Coverage

6.5 Geophone Spread Geometries

Module 7: Seismic Noise

7.1 Signals and Noise

7.2 Coherent Noise

7.3 Incoherent Noise

7.4 Aliased Frequencies

7.5 Multiples

Module 8: Seismic Data Processing I

8.1 Propagation of Seismic Waves through Earth

8.2 Mechanical Processes

8.3 Interactive Processes

8.4 Spherical Divergence Compensation and Gains

8.5 Band Pass Filter

8.6 Deconvolution

Module 9: Seismic Data Processing II

9.1 Seismic Data Processing Flow

9.2 Dynamic Corrections

9.3 Velocity Analysis

9.4 Residual Statics

9.5 Migration

Module 10: Seismic Resolution

10.1 Resolution

10.2 Seismic Resolution

10.3 Fresnel Zone

10.4 More on Seismic Resolution

Module 11: Seismic Interpretation

11.1 Seismic Data Display Standards

11.2 Seismic Section Display Scales

11.3 Base Map

11.4 Seismic Interpretation

11.5 Time to Depth Conversion

11.6 2D Seismic Modeling

11.7 Synthetic Seismogram

Module 1

Physics and

Electromagnetic Spectrum

At the end of this module you would be

able to understand

� Science behind Imaging the

Invisible

� Basic Methodology of Geophysical

Exploration Techniques

1.1 Basic Foundations of Physics

Physics is the science of Matter and Energy. Mater and Energy are related to

each other through Einstein’s equation:

E = m c2

All Matter-Energy in the Universe has a dual nature. They exist as Particles

as well as Waves. Thus Physics is the science of Particles and Waves

The first question arises – What is the Size of Particles ?

There is a wide range of particle size, considering the universe itself as a

particle down to the elementary sub-atomic particles. Thus physics is

divided into different branches on the basis of size of the particles as shown

below.

Accordingly Geophysics is the physics of Earth, Space and Planets. It

includes the study of solid earth, its interior, fluid envelopes (oceans) and

atmospheres.

The second question arises – What is the Frequency of Waves ?

This includes the whole electromagnetic spectrum (EMS). If we extend

down the EMS to sound and ultra-sound waves then we have a whole range

of frequencies that exist in the universe as shown below.

If we consider our human sensors (eyes and ears), we can only receive the

sound and visible light frequencies, thus only a narrow window in the EMS

is Visible to us; the rest is totally invisible as shown.

Sound/Ultra Sounds

101

Visible Invisible Invisible

1.2 Imaging Principle

The imaging principle considers three components; an energy source in the

form of some radiation having a band of frequencies, a set of mediums

through which the radiation passes and a sensor which can receive the given

frequencies.

This principle holds true in case of our eyes. We need a light source which

transmits frequencies in the visible band. These frequencies hit the surfaces

of different materials and are reflected back and finally received by our eyes

which act as sensors. This produces the sensation of vision. Each material

based on its physical properties (albedo) absorbs certain frequencies and

reflects the remaining frequencies, which generates the impression of color

on our retina.

Similarly bats have no eyes, instead they use another band of frequencies,

ultra-sounds to get vision.

Thus different parts of the electromagnetic spectrum can be used for various

types of imaging. Though most of the EMS is invisible to us, but special

tools have been developed which have sensors that can receive a certain

band of frequencies. The received radiation can be processed and translated

into some graphical form which can be viewed and interpreted by human

eye as illustrated below.

This imaging the invisible is used in medical instrumentation, geophysical

methods, astrophysics and various security scanners.

Sensor

Translator

Display

1.3 Imaging the Invisible

Some common imaging the invisible instruments or techniques are

summarized below.

• Medical Instrumentation

- X-rays

- Ultra Sound

- Magnetic resonance Imaging (MRI)

- Computerized Tomography Scan (CT-Scan)

• Scanning Tunnel Microscope

• Astrophysics

- Radio Telescope

- Infrared telescope

• Geophysics

- Gravity & Magnetic

- Electrical & Electromagnetic

- Ground Penetration radar (GPR)

- Seismic

1.4 Fundamental Laws of Wave Propagation

Each imaging method is based on some contrast in a physical parameter. The

physical parameter represents a potential field and therefore occupies a

portion of the Electro Magnetic Spectrum (EMS).

As each method uses a different frequency band of EMS, two generalized

laws can be defined which can be applied to all imaging techniques.

Frequency verses Resolution

Resolution of a system or technique is defined as its ability to view the

smallest size of an object. This purely depends on the frequency used by the

system. We know the wavelength is inversely proportional to frequency.

According the smallest size that can be viewed using a frequency is 1/4th of

the wavelength of that frequency. Thus higher the frequency the higher

would be the resolution.

- High Frequency > Small Wavelength > View Small Objects

- Low Frequency > Large Wavelength > View Large Objects

Frequency versus Penetration

Another important concern is the wave penetration or imaging depth. In this

regard the frequency is inversely proportional to depth of penetration. Thus

higher the frequency the lower would be the penetration.

- High Frequency > Less Penetration

- Low Frequency > Deep Penetration

Thus increasing the frequency increases the resolution but decreases the

depth of penetration.

Considering the above two principles various imaging methods have been

devised according to their application and usage. In this regard, just consider

the following two examples.

Seismic

Seismic methods use low frequencies (10-200 Hz). Thus they have low

resolution but high depth of penetration. This suits us for imaging the earth.

The thinnest layers are in order of several meters and may be few kilometers

deep. The seismic waves can propagate down to such depths and resolve

these layers.

Ultrasound

Ultrasound uses comparatively higher frequencies (> 20 KHz). This

increases the resolution to millimeters but decreases the depth of penetration

to less than a meter. It can be used successfully to image small tissues in the

human body. Thus there is no need for deeper penetration.

Module 2

Electrical Resistivity

Methods

At the end of this module you would be

able to understand

� Resistivity Meter

� Schlumberger and Wenner

Configuration

� Vertical Electrical Sounding

Curves

2.1 Electrical Resistivity Methods

The electrical resistivity method is based on resistivity (opposite of

conductivity) contrast. Thus it involves measuring the apparent resistivity of

soils and rock as a function of depth or position. The resistivity of soils is a

complicated function of porosity, permeability, ionic content of the pore

fluids, and clay mineralization. The unit of resistivity is ohm-meter.

The most common electrical methods used in mineral exploration,

hydrogeologic and environmental investigations are vertical electrical

soundings (VES) (resistivity soundings) and resistivity profiling.

The VES techniques are used to determine depth to groundwater, map clay

aquitards, saltwater intrusion and vertical extent of certain types of soil and

groundwater contamination, characterize subsurface hydrogeology,

determine depth to bedrock/overburden thickness, map stratigraphy and

estimate landfill thickness

The resistivity profiling techniques are used to map lateral extent of

conductive contaminant plumes, explore for sand and gravel and delineate

disposal areas.

Resistivities of some common rocks and mineral are given below in ohm-

meter.

- Igneous and Metamorphic Rocks

Granite 5x103 – 10

6

Basalt 103 – 10

6

Slate 6x102 - 4x10

7

Marble 102 - 2.5x10

8

Quartzite 102 - 2x10

8

- Sedimentary Rocks

Sandstone 8 - 4x103

Shale 20 - 2x103

Limestone 50 - 4x102

- Soils and waters

Clay 1 - 100

Alluvium 10 - 800

Groundwater (fresh) 10 - 100

Sea water 0.2

2.2 Resistivity Meter

The instrument used to carry out electrical resistivity surveys is called

resistivity meter. It consists of the following two main units:

Transmitter

It includes the battery and an ammeter and sends out well defined regulated

current to the ground through the current electrodes. The current can be

direct current or low frequency alternating current.

Receiver

The receiver consists of a voltmeter and detects the transmitted signal

current by measuring the potential developed between the two potential

electrodes.

Modern digital instruments also contain an analog to digital converter and a

microprocessor which quickly takes multiple readings and averages them to

get reliable results.

The working of a resistivity meter is shown in the following figure. Here C1

and C2 are current electrodes, P1 and P2 are potential electrodes, A is

ammeter and V is voltmeter.

2.3 Electrical Resistivity Surveying

During a resistivity survey, current is injected into the earth through a pair of

current electrodes, and the potential difference is measured between a pair of

potential electrodes. The current and potential electrodes are generally

arranged in a linear array. Common arrays include the Wenner array,

Schlumberger array, dipole-dipole array and pole-Dipole array. The apparent

resistivity is the bulk average resistivity of all soils and rock influencing the

current. It is calculated by dividing the measured potential difference by the

input current and multiplying by a geometric factor specific to the array used

and electrode spacing as given below;

VR k

I

∆=

where ∆V is the potential difference, I is the current and k is a geometric

factor depending on the geometry of the array.

In vertical electrical soundings, the distance between the current electrodes

and the potential electrodes is systematically increased, thereby yielding

information on subsurface resistivity from successively greater depths. The

variation of resistivity with depth is modeled using forward and inverse

modeling computer software. Thus this technique provides a 1D vertical

model of the subsurface.

In resistivity profiling, the electrode spacing is fixed and measurements are

taken at successive intervals by moving the entire array along a profile. This

gives some information about lateral changes in the subsurface resistivity,

but it cannot detect vertical changes in the resistivity. Data are generally

presented as cross-section profiles or contour maps and interpreted

qualitatively.

2.4 Electrode Geometries

Some common electrode geometries are illustrated in the next figure along

with their geometric factors. The depth of penetration of most of these

configurations is half the geometry spread length.

Among these geometries Schlumberger and Wenner are the most widely

used configurations. Each geometry has advantages and disadvantages.

Advantages of Wenner array as compared to Schlumberger array include;

large potential electrode spacing places less demand on instrument

sensitivity and simplicity in geometric factor equation due to equally spaced

electrodes. The main disadvantages of Wenner array are that in an

expanding array all electrodes must be moved for each reading which is not

the case with Schlumberger array and secondly it is more sensitive to local

near-surface lateral variations.

2.5 Resistivity Interpretation

Resistivity data is interpreted in the following way;

Qualitative Interpretation

In this type of interpretation the apparent resistivity values are directly used.

For a random or gridded distribution of resistivity stations, iso-resistivity

contour maps are generated for a particular depth to show the general

distribution of resistivity at the given depths. Similarly resistivity stations

are joined along a profile and cross-sections of apparent resistivity are

generated. Such sections are also generated for resistivity profiling

techniques. These sections show the cross-sectional variation of resistivity as

illustrated in the next figure.

Quantitative Interpretation

The main task in quantitative interpretation is to identify the subsurface

layers and get their true resistivities, which in turn are translated into

geological formations. The measured apparent resistivity values for 1D VES

are normally plotted on a log-log graph paper. To interpret the data from

such a survey, it is normally assumed that the subsurface consists of

horizontal layers. In this case, the subsurface resistivity changes only with

depth, but does not change in the horizontal direction. The true resistivities

are obtained through a manual procedure of matching segments of field

resistivity curves with a set of 2 layer master. Two layer master curves are

illustrated below.

Several reverse and forward modeling techniques are also available that can

interpret the apparent resistivity field curves into true resistivities and depths

data as illustrated in the next figure.

Module 3

Gravity & Magnetic

Methods

At the end of this module you would be

able to understand

� Gravimeter & Magnetometer

� Gravity Field Correction: Free Air

and Bouguer Anomaly

� Terrain Corrections

� Regional and Residual Separation

� Gravity Modeling

� Magnetic Data Processing

� 2D Grid Processing

3.1 Gravity and Magnetic Prospecting

Gravity and magnetic prospecting techniques involves measuring passive

potential fields of the Earth. In both these methods the measured signal is a

composite contribution from all depths. On the other hand, seismic

prospecting can give a detailed picture of Earth structure with different

subsurface components resolved. Thus seismic method has much higher

resolution as compared to these methods.

Gravity and magnetic methods can be carried out on land or sea using

different techniques and equipment. In addition aero-gravity and magnetic

surveys can also be conducted.

3.2 Gravity Method

In all gravity surveys the vertical component of g is measured. Gravity

prospecting can be used where density contrasts are present in a geological

structure, and the usual approach is to measure differences in gravity from

place to place. In gravity prospecting we are mostly interested in lateral

variations in Earth structure which in turn create lateral variations in density.

Gravity method was first applied for prospecting salt domes in the Gulf of

Mexico, and later for looking anticlines in continental areas. Gravity cannot

detect oil directly, but if the oil is of low density and accumulated in a trap,

it can give a gravity low that can be detected by gravity prospecting.

Anticlines can also give gravity anomalies as they cause high or low density

beds to be brought closer to the surface. Nowadays in the petroleum

industry, gravity method is used for regional studies to identify large and

thick enough sedimentary basins, as sedimentary rocks have lower densities

than basement rocks. Gravity prospecting can also be used for mineral

exploration if substantial density contrasts are expected, such as, chromite

bodies have very high densities, buried channels which may contain gold or

uranium can be detected because they have relatively low density.

The unit of gravity is Gal, after Galileo, where 1 Gal = 1 cm/sec2.

Thus g at

the surface of the Earth is approximately 103 Gals. Gravity anomalies are

measured in units of milliGals, where 1 mGal = 10-3

Gals = 10-5 m/sec2.

The densities of few common minerals in g/cm2 are given below

- Quartz 2.65

- Felspar 2.6

- Biotite mica 2.9

- Calcite 2.6 – 2.7

3.3 Gravimeters

Gravity meters, usually called gravimeters, are sensitive to 0.01 mGal = 10-8

of the Earth’s total value. Thus the specifications of gravimeters are amongst

the most difficult to meet in any measuring device. It would be impossible to

get the accuracy required in absolute gravity measurements quickly with any

device, and thus field gravity surveying is done using relative gravimeters.

There are two basic types of gravimeters:

Stable Gravimeters

These work on the principle of a force balancing the force of gravity on a

mass, such as the Gulf gravimeter. These gravimeters take a long time to

measure each point. The Gulf gravimeter comprises a flat spring wound in a

helix, with a weight suspended from the lower end. An increase in g causes

the mass to lower and rotate. A mirror on the mass thus rotates and it is this

rotation that is measured. The sensitivity of these gravimeters is ~ 0.1 mGal.

They are now obsolete, but a lot of data exist that were measured with such

instruments and it is important to know that such data are not as accurate as

data gathered with more modern instruments

Unstable Gravimeters

These are virtually universally used now. They are well devised mechanical

devices where increase in g causes extension of a spring, but the extension is

magnified by mechanical geometry. An example is the Wordon gravimeter,

which has a sensitivity of 0.01 mGal, and is quite commonly used. Wordon

gravimeter is shown in the next figure. It is housed in a thermos flask for

temperature stability, but it also incorporates a mechanical temperature

compensation device. It is evacuated to eliminate errors due to changes in

barometric pressure. It weighs about 3 kg and the mass weighs 5 mg.

Vertical movement of the mass causes rotation of a beam, and equilibrium is

restored by increasing the tension of torsion fibers.

Another commonly used gravity instrument is LaCost-Romberg gravimeter.

The latest gravimeters are completely electronic with a software controlled

interface and a built-in GPS. They directly store data on a media floppy,

which can be downloaded to a computer for further processing.

3.4 Gravity Surveying and Corrections

Gravity field procedure involves measurement on a base station followed by

measurements on a number of stations and finally repeating the base station.

For larger surveys base station must be repeated approximately every two

hours.

During the survey, at each station the following information is recorded:

- Time at which the measurement is taken.

- Reading of the Gravity Meter in scale readings

- Navigation Data: Latitude, longitude and elevation of the station.

A set of corrections are applied to the observed gravity data which are

discussed below.

Instrument Calibration

Each instrument has a scale constant (SC), provided by the manufacturer,

that translates scale readings (SR) into mGal as given below.

gobs = SR * SC

Drift Correction

The drift correction incorporates the effects of instrument drift,

uncompensated temperature effects and the gravitational attraction of the sun

and moon. It is computed by taking two reading at the base station, one at

start and the other at end of survey. The drift rate is computed as;

_ _

_ _

base start base end

base end base start

g gDR

t t

−=

−

Now drift correction for a station is given by;

_*( )stat base start

DC DR t t= −

Latitude Correction

This correction is needed because of the ellipticity of Earth as g is reduced at

low latitudes because of the Earth’s shape and rotation. It is given by;

.0008122 2NS base

LC Dist Sin φ=

where DistNS is North-South Distance, φbase is latitude of base. In N-hemisphere this correction is negative if station is towards north of

base and positive if station is south of base and vise versa for southern

hemisphere.

Free Air Correction

The correction is also called elevation correction. It is required to correct for

the variable heights of the stations above sea level, because g falls off with

height. It is given by;

( )stat base

FAC k E E= −

where Estat is elevation of station, Ebase is elevation of base and

k = .9406 for feet and k = .3086 for meters.

Bouguer Correction

This correction accounts for the mass of rock between the station and sea

level. It has the effect of increasing g at the station, and thus it is subtracted.

Bouguer correction is given by;

( )stat base

BC k E Eρ= −

where Estat is elevation of station, Ebase is elevation of base, ρ is density of the

material and k = .01276 for feet and k = .04185 for meters.

Bouguer Anomaly

The Bouguer anomaly is computed by apply all the above corrections to the

observed gravity as given by;

_( )stat base start

BA g g DC LC FAC BC= − + + + −

Terrain Corrections

The effect of terrain always reduces the observed g. This is true for a

mountain above or a valley below the station, both cause g to be reduced.

Previously terrain corrections were done by hand using a transparent

graticule (shown below) placed at the station, then average height of each

compartment is estimated and Hammer chart was used to obtain the

correction for the station. This chart gives the correction for a particular

distance from the station. It has been worked out assuming a block of

constant height for each compartment. This manual procedure was very

time-consuming and involved a lot of repetition. With the availability of

digital terrain models the same procedure has been computerized.

The figure below shows an observed gravity anomaly.

The following are the drift, latitude, free air and Bouguer corrections for the

above observed anomaly.

The observed anomaly after application of above corrections is shown

below.

3.5 Regional Residual Separations

The processed gravity anomaly contains regional and residual effects. For

regional studies we are interested in regional anomaly while for local studies

we are interested in residual anomaly. There are several techniques for

separation of regional and residual trend such as graphical method, moving

average method with 3, 5 or 7 points operator and statistical best fir or

regression techniques from first to higher orders. The figure below shows

processed anomaly along with its regional and residual components.

3.6 Gravity Modeling

The observed gravity anomaly gives us the trend of gravity along a profile.

Forward modeling techniques, such as the Talwani method are used to create

a subsurface model of geological bodies each assigned with a density. The

modeling process generates a model anomaly. The shape and density of

subsurface bodies are changed in such a way that the model anomaly

matches the observed anomaly. Previously this was done manually through a

tedious process. Currently interactive applications are available that can

quickly create and fit the model.

The next figure shows a subsurface regional model with its model curve

fitted to the regional trend of gravity.

3.7 Magnetic Method

In magnetic prospecting precise magnetic field is measured to locate

geological structures and man-made objects in the ground or under the sea.

The number of possible applications of magnetic exploration is unlimited

and include; oil and gas exploration, mineral exploration such as iron ore,

underground pipeline detection, buried unexploded ordnance detection, and

archeological prospecting. All these objects are detected as they posses an

extremely weak magnetic field of their own, which is a measureable local

disturbance in the Earth’s magnetic field. Such disturbance is called a

magnetic anomaly. This contrast in magnetic field is called susceptibility. It

is measurement in Gammas.

3.8 Magnetometers

A device that measures the magnetic fields is called a magnetometer. There

are several types of magnetometers among which the two most common

types are;

Proton Precession Magnetometer

Proton precession magnetometers, also known as proton magnetometers,

measure the resonance frequency of protons (hydrogen nuclei) in the

magnetic field to be measured, due to nuclear magnetic resonance (NMR).

As the precession frequency depends only on atomic constants and the

strength of the ambient magnetic field, the accuracy of this type of

magnetometer is very good thus it is widely used in magnetic prospecting.

This magnetometer measure the total intensity (T) of the magnetic field.

Fluxgate Magnetometer

A fluxgate magnetometer consists of a small, magnetically susceptible, core

wrapped by two coils of wire. An alternating electrical current is passed

through one coil, driving the core through an alternating cycle of magnetic

saturation; i.e., magnetised, unmagnetised, inversely magnetised,

unmagnetised, magnetised, etc. This constantly changing field induces an

electrical current in the second coil, and this output current is measured by a

detector. In a magnetically neutral background, the input and output currents

will match. However, when the core is exposed to a background field, it will

be more easily saturated in alignment with that field and less easily saturated

in opposition to it. Hence the alternating magnetic field, and the induced

output current, will be out of step with the input current. The extent to which

the input and output currents are out of step, will depend on the strength of

the background magnetic field. Often, the current in the output coil is

integrated, yielding an output analog voltage, proportional to the magnetic

field. This magnetometer measures the vertical component (Z) of the

magnetic field.

3.9 Magnetic Surveying and Corrections

The processing of observed magnetic data is similar to gravity data, except

the types of corrections are totally different. Only two corrections are

applied to magnetic data which are discussed below.

Diurnal Correction

The suns solar activity continuously disturbs the Earth’s magnetic field.

Thus we need to remove these effects from the observed data. These effects

can be removed in two ways. With a single magnetometer the procedure is

similar to drift correction. We start with a base station and repeat it at the

end and from the difference in the two readings we compute a drift rate

which is applied to observations of all stations. On the other hand if we have

two instruments we fix one at the base while the other takes readings at the

stations. Later the readings at the base are used as corrections that are

applied to the base stations at the corresponding times.

Normal Correction

The Earth’s magnetic field is not constant and changes with latitude and

longitude. Global magnetic anomaly maps are published annually. They are

used to apply corrections at the station locations.

3.10 2D Grid Processing

In the above sections we discussed processing of gravity and magnetic data

along a profile. If there is a set of parallel profiles they make up a grid. The

data corrections or reductions are applied individually along a profile, but

regional residual effects are separated through grid processing. In these

techniques a 3 x 3 operator moves step by step through the grid to compute

regional anomaly at the center of the operator. This technique is similar to

Griffin’s method which used a circular operator around the grid node. The

computed regional trend grid is subtracted from the input observed gravity

grid to get a residual anomaly grid as shown in the next figure.

Module 4

Seismic Waves and Rock

Physics

At the end of this module you would be

able to understand

� Types of Seismic Waves

� Stress and Strain

� Hook’s Law of Elasticity

� Seismic Velocities

� Engineering Properties and Rock

Physics

4.1 Seismic Waves

Seismic waves travel through the Earth, as the result of a tectonic earthquake

or an explosion. They propagate through a medium similar to sound waves.

They are also called Elastic Waves.

4.2 Types of Seismic Waves

Seismic waves are classified into the following types:

• Body Waves

- P-Waves

- S-Waves

• Surface Waves

- Rayleigh Waves

- Love Waves

Body Waves

Body waves travel through the interior of the Earth. They follow ray-paths

bent by the varying density and modulus (stiffness) of the Earth's interior.

Body waves are further classified into P and S waves.

P-Waves P means Primary Waves as they are fast and arrive first. They are also

called longitudinal or compressional waves, as particle motion is parallel to

wave propagation. The ground is alternately compressed and dilated in the

direction of propagation. These waves can travel through any type of

material. P-waves have the following velocities in different mediums.

- Air 330 m/s (Take the form of sound waves, thus travel at the speed of sound)

- Water 1450 m/s

- Granite 5000 m/s (In Solids twice as fast as S Waves)

When generated by an earthquake they are less destructive than the S waves

and surface waves that follow them, due to their smaller amplitudes

P Wave propagation is analogous to sound waves as shown.

S-Waves S means Secondary Waves as they arrive after the P Waves. They are also

called transverse or shear waves, as particle motion is perpendicular to wave

propagation. The ground is displaced perpendicularly to the direction of

propagation. In the case of horizontally polarized S waves, the ground

moves alternately to one side and then the other. They travel only through

solids, as fluids (liquids and gases) do not support shear stresses

Their speed is about 60% of that of P waves in a given material

S waves are several times larger in amplitude than P waves for earthquake

sources.

S wave propagation is analogous to Light as shown below.

Surface Waves

Surface waves are analogous to water waves and travel just under the Earth's

surface. They travel more slowly than body waves. Because of their low

frequency, long duration, and large amplitude, they can be the most

destructive type of seismic wave.

Rayleigh Waves They travel only under the Earth’s surface. They are in the form of ripples,

similar to those on the surface of water. Rayleigh waves are also called

ground roll in seismic exploration data. Their speed is about 70% of that of S

waves. Their existence was predicted by John William Strutt, Lord Rayleigh,

in 1885.

Rayleigh wave propagation is analogous to Ocean surface as shown below.

Love Waves They travel only under the Earth’s surface. They cause horizontal shearing

of the ground.

Their speed is about 90% of that of S waves, slightly faster than Rayleigh

waves.

Named after A.E.H. Love, a British mathematician who created their

mathematical model in 1911

Love wave propagation is analogous to movement of Snake or Shaken Rope

as shown below.

Seismic energy released during an earthquake can be recorded on a

seismogram as shown below.

4.3 Uses of Seismic Waves

• P Waves are commonly used in Oil & Gas Exploration

• Special 3 component (3C: P, SH, SV) Surveys are also carried out for

rock physics and reservoir analysis.

• P & S Waves are used to determine the engineering properties of ground.

• Earth’s liquid outer core was discovered due to the fact that shear waves

cannot pass through liquids (as demonstrated by Richard Dixon Oldham)

Minutes

Surface waves

0

P S

10 20 30 40 50

‘Primary’ (first to arrive) ‘Secondary’ (second to arrive)

‘Surface’ (last to arrive)

4.4 Stress and Strain

Stress: The force causing the deformation in a material. Stress can be of two

types.

Normal stress is applied perpendicular to the face of material.

Shear stress is applied parallel or tangential to the face of a material.

Strain: The amount by which a material body is deformed. Strain is also of

two types.

Normal strain acts perpendicular to the face of a material that it is acting on.

Shear strain acts parallel to the face of a material that it is acting on.

4.5 Elasticity & Stiffness

Elasticity: A material is said to be elastic if it deforms under stress (applied

force), but returns its original shape when the stress is removed. The amount

of deformation is called the strain.

Stiffness is the resistance of an elastic body to deflection or deformation by

an applied force.

Stress-Strain relation of rock deformation is illustrated in the following

figure.

Each material has an elastic limit and a fracture point. Stress applied within

the elastic limits will cause strain that will be recovered when the stress is

removed. Stress greater than the elastic limit, but below the fracture point

will cause a permanent strain, which will not recover when the stress is

removed. If the stress is increased to or above the fracture point the material

will break up.

On the basis of stress stain analysis, there can be two types of deformation:

Elastic Deformation: A temporary change in shape or size that is recovered

when the applied stress is removed.

Ductile (Plastic) Deformation: A permanent change in shape or size that is

not recovered when the applied stress is removed.

4.6 Hooks Law of Elasticity

The amount by which a material body is deformed (the strain) is linearly

related to the force causing the deformation (the stress)

F = - k x Where

X is the distance the material is stretched or compressed away from

equilibrium position (Meter)

F is the restoring force exerted by the material (Newton)

K is Spring Constant (Newton/Meter)

Hook’s law is illustrated in the following figure.

4.7 Elastic Moduli

Elastic moduli for homogeneous and isotropic materials are discussed

below.

Bulk Modulus

The bulk modulus (K) of a substance measures the substance's resistance to

uniform compression. It is the ratio of volume stress to volume strain. It is

defined as the pressure increase needed to affect a given relative decrease in

volume. It describes the material's response to uniform pressure. For a fluid,

only the bulk modulus is meaningful.

Young’s Modulus

Young's modulus or modulus of elasticity (E) is a measure of the stiffness of

an isotropic elastic material. It is the ratio of the uniaxial stress over the

uniaxial strain in the range of stress in which Hooke's Law holds. It

describes the material's response to linear strain.

Lame’s Constant

The Lame’s Constant (λ) has no physical interpretation, but it serves to

simplify the stiffness matrix in Hooke's law. It is also called Lame’s First

Parameter.

Shear Modulus

Shear modulus or modulus of rigidity (µ), is defined as the ratio of shear

stress to the shear strain (angle of deformation). It is concerned with the

deformation of a solid when it experiences a force parallel to one of its

surfaces while its opposite face experiences an opposing force (such as

friction). It describes the material's response to shearing strains.

Poisson’s Ratio

Poisson's ratio (σ) is the ratio of transverse strain (normal to the applied

load) to longitudinal strain (in the direction of the applied load). When a

sample of material is stretched in one direction, it tends to contract (or

rarely, expand) in the other two directions. Conversely, when a sample of

material is compressed in one direction, it tends to expand (or rarely,

contract) in the other two directions. Poisson's ratio is a measure of this

tendency.

P-Wave Modulus

P-wave modulus (M) or longitudinal modulus is the ratio of axial stress to

axial strain in a uniaxial strain state.

4.8 Computing Density/Moduli from Seismic Velocities

A seismic survey provides velocity information about sub-surface layers.

Once the P or S wave velocity of a material is determined, its density and all

moduli can be computed. This determination of such parameters is termed as

Rock Physics or Engineering Properties.

A set of rock physics equations are listed below.

P-Wave Velocity

S-Wave Velocity

Density

Vp Vs Ratio

Bulk Modulus

Young’s Modulus

Lame’s Constant

Shear Modulus

Poisson’s Ratio

P Wave Modulus

1 16 1 36p s

V . *V .= +

( 1 36) /1.16s pV V .= −

.250.31*p

Vρ =

4

3Ratio

KVpVs

µ= +

2 243

( )p sK V Vρ= −

9

3

KE

K

µ

µ=

+

2

3K

µλ = −

2

sVµ ρ=

2 2 2 20.5( 2 ) / ( )p s p sV V V Vσ = − −

4

3M K

µ= +

Module 5

Seismic Refraction Methods

At the end of this module you would be

able to understand

� First Breaks: Direct and Refracted

Waves

� Automated First Break Picking

� TX-Graphs, Layer Velocity and

Depth Computation

� Low Velocity Weathered Layers

and Statics Computation

5.1 Snell’s Law

Snell’s law was originally developed for light waves, but it can be equally

applied to sound and seismic waves. Accordingly when a wave enters from a

less dense medium (ρ1) to a denser medium ((ρ2), it bends away from the

normal. Thus the angle of incidence (i) is less than angle of refraction (r). If

we keep on increasing angle i, angle r will also increase, until it becomes

90º. The angle i for which angle r is 90º is called critical angle (ic).If angle i

become greater than ic , the wave is reflected back into the same medium. In

case of seismic the angle i increases with offset, the distance between source

and receiver. Thus on the basis of angle i seismic wave is split into three

components as shown below. The transmitted wave acts as a secondary

source and is again split into three components at the next interface. This

continues as the waves move down into deeper layers and forms the basis of

seismic refraction and reflection methods.

It must be noted that the transmission, refraction and reflection of seismic

waves only takes place when the velocity of each underlying layer is higher

than that of above it. In Earth the general trend is increase of velocity with

depth. As we move down towards deeper layers the overburden pressure

increases, which increases the density and hence the velocity also increases.

ρ1

ρ2

ρ2 > ρ2

Interface

Transmission

Refraction

Reflection

i < ic

i = ic

i > ic

5.2 Seismic Refraction Method

After gravity method, seismic refraction method was developed for

exploration of hydrocarbons. In 1924 it was first used for delineation of

shallow salt domes. Due to some limitations it was soon replaced with

seismic reflection method, which continues to be the main technique for

exploration of hydrocarbons and imaging deep structures. Today the seismic

refraction method is widely used to delineate the Low-Velocity Weathered

Layers for computation of Statics corrections that are applied to the main

reflection data. In addition it is considered as a valuable tool for near-surface

geophysics & engineering, such as delineation of bed rock or basement and

determining the engineering properties of ground. In the following sections

the complete workflow of refraction method is discussed.

5.3 Seismic Refraction Data Acquisition

A seismic refraction recorder usually consists of 24 channels each of which

is connected to a geophone. The geophones are placed along a profile with

variable geophone intervals. Two shots are taken at both ends of the profile,

the first near geophone # 1 is called forward shooting while the second near

geophone # 24 is called reverse shooting. This results in two seismic

monitors each with 24 seismograms (traces) as shown below.

1 24

Forward Shooting Reverse Shooting

5.4 First Breaks

First breaks are the events that reach first at a geophone and are also called

First Arrivals. They are the first prominent wave amplitude on a seismogram

(trace) as shown below.

In seismic refraction techniques, we need to pick the first breaks times from

the seismic traces. This time can be picked in four different modes as shown.

First Breaks must be picked in any one of the four modes, but the selected

mode must be used for the whole project. It must be noted that within a

project two different modes cannot be used.

Initially the seismic monitors were in the form of papers, thus first breaks

were marked with a pen and their arrival times were noted. This is referred a

hand picking. With the advent of digital data and computers, interactive

software became available which provided a computer aided environment

for picking first breaks using the mouse. This is referred as manual picking.

With the increased usage of artificial intelligence in geosciences, several

neural-network based techniques have been developed for automated

picking of first breaks.

Zero Crossing Positive Slope Crest

Zero Crossing Negative Slope Trough

A seismic monitor with artificial intelligence based first break picks is

shown below.

Let’s consider a three layer earth model with velocities Vo, V1 and V2

respectively and a seismic refraction spread as shown below. Now at near

offset geophones the direct waves representing Vo .reach first, then at the

next few geophones the refracted waves from top of V1 reach first and

finally at the far offset geophones refracted waves from top of V2 reach first.

The point at which the direct and refracted waves reach at the same time is

called crossover distance.

ρ0 V0

ρ1 V1

ρ2 V2

Direct Waves

Refracted Waves

Refracted Waves Crossover Distance

Distance

Time x1,t1

x2,t2

S = dt/dx = (t2-t1)/(x2-x1)

V = 1/S = (x2-x1)/(t2-t1)

5.5 Time-Distance Graphs

The picked first break arrival times, for both forward and reverse shooting,

are plotted on a graph paper against their offsets as shown below.

For both forward and reverse shooting data, best-fit lines are passed through

each segment of data, representing a subsurface layer as shown below.

Distance (X) Distance (X) Forward Shooting Reverse

Shooting

Time

Cross Over Distance

Intercept Time

Total Time

Vo Vo

V1 V1

Time

The velocities of best-fit lines (Vo, V1) are computed from their respective

slope. Similarly the crossover distance (Xc), intercept time (Ti) and total time

(Tt) are marked on the graph as shown on the pervious figure.

The thickness of the first layer (Ho) can be computed by any one of the

following equations. The first uses intercept time while the other uses

crossover distance.

Thus using the refraction method the velocities and thicknesses of near

surface weathered layers are determined which are used for computation of

statics corrections, for application to seismic reflection data.

5.6 Statics Corrections

The Earth’s near surface is made up of weathered and sub-weathered low-

velocity layers composed of unconsolidated material. These layers induce a

delay in seismic reflection data events which distort the continuity in

subsurface layers geometry as seen in seismic sections.

Statics are corrections applied to seismic reflection data to remove the effect

of weathered layers and elevation. These corrections are applied by reducing

the data with respect to a Datum Plane.

The following figure illustrates the effect of weathered layers in imaging a

horizontal sub-surface reflector on a seismic section. Due to variations in

weather layer thickness and/or velocity, each recorded trace experiences a

different delay time. Thus the horizontal reflector attains the shape of the

weathered layer.

1

2 2

12

oo

o

V VTiH

V V=

−

1

12

oo

o

V VXcH

V V

−=

+

Similarly topographic (elevation) variations also affect the shape of the sub-

surface reflectors in seismic data, due to variable wave travel paths. Thus the

horizontal sub-surface reflector appears as an inverted image of the

topography on a seismic section as shown below.

Low Velocity Layer

Reflector

More Delay

Less Delay More Delay

Reflector

Longer Path Shorter

Path

Longer Path

Topography

From the above discussion it is clear that statics corrections must be

computed in order to remove the effects of weathered layers and topography

as shown below.

It can be seen from the above equation that statics correction has two

components; weathered layer and elevation statics. For the weathered layer

statics velocity and thickness of weathered layers is provided by the

refraction method. For the elevation (E) statics, we get the elevation from

navigation data, the replacement velocity (VR) is selected above the highest

sub-weathered velocity. Its value can be selected somewhere above 2000

m/sec and it must be kept constant throughout the project. The selection of

datum (D) is arbitrary; it can be selected above or below the weathered

layers. The figures below show a seismic section without (left) and with

(right) statics corrections.

11

1 1

[ ... ]*1000o n

o n R

H HH E DStatics

V V V V

−

−

−= − + + + +

Reflector

Topography

Weathered

Sub-Weathered

Vo

V1

Datum

Ho/Vo

H1/V1

(Elevation - Datum) / Replacement Velocity

1

1

[ ]*1000n

i

i i R

H E DStatics

V V

−

=

−= − +∑

Weathered Layer Statics

Elevation Statics

5.7 Limitations of Seismic Refraction Method

Some limitations of refraction method are summarized below:

The refraction method requires the velocity of sub-surface layers must

increase with depth. If a low velocity layer under a high velocity layer is

encountered, it cannot be detected by refraction method.

Similarly blind zones cannot be detected by this method. If thickness of a

layer is less as compared to the layers above and below it and/or the velocity

contrast between it and the layer that underlies it is inadequate, such layer

cannot be delineated by this method.

Refraction method requires a larger spread length as compared to reflection

method, for mapping the same interface (at the same depth). Thus for

imaging deeper layers refraction method needs a much longer spread, which

makes it impractical for such applications.

Module 6

Seismic Reflection Data

Acquisition

At the end of this module you would be

able to understand

� Digital Sampling

� Seismic Sensors: Geophones as

Transducers

� Seismic Recorder and

Multiplexing

� Seismic Sources: Dynamite and

Vibroseis

� Spread Geometries & Fold

Coverage

6.1 Digital Sampling and Aliasing

The initial seismic instruments recorded data on a moving paper, the next

generation instruments recorded analog seismic data on magnetic tapes.

With the advent of computers and digital systems, the seismic data was

recorded in digital form in some standard format. Today all processing and

interpretation is performed on digital seismic data, thus it is necessary to

understand the difference between analog and digital data and the sampling

theory.

Let’s consider an analog representation of a sinusoidal function as shown

below. In this form the variation of amplitude with time is recorded

continuously.

Now the same sinusoidal function can be represented in digital form as

shown below. It can be seen that discrete samples of amplitude values are

taken after fixed interval of time called sampling interval. Joining these

samples the sinusoidal function can be reconstructed.

If �t is in milliseconds the Sampling frequency i.e. the number of samples

per second, is given by;

Now the Nyquist frequency, which is the highest recoverable signal

frequency for the given sampling frequency is given by;

Thus for a given sampling interval, the recorded signal frequencies must not

be greater than Nyquist frequency, otherwise they will appear as low

frequencies. This is called Aliasing Effect, caused by coarse sampling

(under-sampling), as shown below.

1000S

ft

=∆

1000

2 2

SN

ff

t= =

∆

The high signal frequencies that are greater than the Nyquist frequency

appear as low frequencies are called Alias frequencies given by;

In seismic recording systems an anti-aliasing filter (fA) having a high-cut

frequency equal to half of Nyquist Frequency is applied to avoid aliasing

effect.

6.2 Seismic Recorder

Seismic Recorder picks seismic signals (vibrations) from geophone sensors

and records them on magnetic media in a digital format. It consists of

multiple channels, each connected to a geophone group. It is similar to an

audio tape recorder which picks audio signals from microphone and records

them on a magnetic tape (cassette).

A geophone is a transducer which transforms mechanical energy (seismic

vibration) into electrical energy. It consists of a moving coil and a stationary

magnet. The movement of the coil due to vibration creates electromagnetic

flux proportional to the magnitude of vibration.

The block diagram of a digital seismic recorder is given below followed by

description of all main modules.

2a N Signal

f f f= −

2

NA

ff =

Preamplifier receives weak signals from geophones and amplifies them.

Filters pass a certain range of frequencies and attenuate the rest. A seismic

recorder has three types of filters.

- Low Cut filter attenuates all frequencies below the cutoff frequency and

passes all frequencies above it. In a seismic recorder it is set to 8-12 Hz

to remove low frequency surface waves called ground roll.

- Notch filter removes only a single frequency and passes all other

frequencies. In seismic it is used to remove 50/60 Hz electric power lines

induction.

- High Cut filter attenuates all frequencies above the cutoff frequency and

passes all frequencies below it. In a recorder it is called Antialias filter,

so that all unwanted high frequencies (above 125 Hz) are removed to

avoid aliasing during analog to digital conversion.

Amplifier further amplifies the signal after filtering unwanted frequencies.

It enhances the gain level of the signal in decibels (db). It must be noted that

the signals received from the geophone and subsequently passing through

preamplifier to amplifier stages are in analog form. After final amplification

the signals will be digitized.

Multiplexer: A seismic recorder consists of several channels (24 to several

hundreds). Each channel is connected to a geophone and comprises of

preamplifier to amplifier stages. Now the signals from all channels need to

be digitized, but the recorder has a single analog to digital converter. Thus a

multiplexer is used to switch one channel at a time. If a recorder has 100

channels and sampling rate is set to 2 milliseconds then the multiplexer must

switch 100 channels within 2 milliseconds to get the first sample of each

channel and be ready to get the next samples. The time available to scan a

channel is called Skew rate given by;

Skew Rate = Sampling Rate / Number of Channels

The working of a multiplexer for a four channel recorder (channels labeled

as A to D) is shown below. Note the multiplexer connects one channel at a

time to the analog to digital converter (ADC).

Analog to Digital Converter (ADC) converts the analog signals into digital

form according to the specified sampling interval. The digital signal consists

of discrete samples represented by a series of amplitude values as a function

of time, called time series.

Formatter/Writer arranges the data samples according to an industry

standard format and writes it to a storage media such as tape, cartridge or

DVD.

A

B

C

D

Time

Geophone

Filters / Amp

ADC / Write Geophone

Filters / Amp

Geophone

Filters / Amp

Geophone

Filters / Amp

B1 A1 C1 D1 A2 SC B2 C2 D2 SC A3 B3 C3 D3 SC …

Multiplexed Data

SC is Scan Code after each time slice.

6.3 Seismic Sources

A seismic source releases energy in the form of elastic waves which

propagate through the earth’s medium. This energy has an amplitude and

phase over a frequency band.

Various types of seismic sources are classified below:

• Land Sources: Used in Land Surveys

- Dynamite: Several Kilograms of dynamite used to generate a short

duration, high energy impulse containing a wide range of frequencies.

Dynamite based seismic data is minimum phase.

- Vibroseis: A truck with a base plate driven by a hydraulic system to

generate a long duration, low energy sweep of defined frequency range.

The vibroseis data needs to be correlated with the pilot sweep. Vibroseis

based seismic data is zero phase.

- Buried Primacord: Explosive extruded into rope-like form having

length of several 100 ft and plowed into ground at 2-3 ft depth. When

detonated at one end or center, the explosive disturbance propagates at

22,000 ft/sec, much higher than seismic velocity in near surface layer.

- In addition there are several other land seismic sources such as weight

dropping, hammer, wooden log hit from side to generate S waves.

• Marine Sources: Used in Offshore Surveys

Explosive Sources using Dynamite

- Flexotir: Small pellet of dynamite embedded in a plastic cartridge. This

charge is detonated at the center of a cast-iron spherical shell towed

behind the ship at 40 ft depth. It pumps out water under high pressure.

- MaxiPulse: Charge packed in a can, injected into the water at 40 ft depth

by a delivery device trailed from the ship. On detonation it forms a

bubble.

Non-Explosive Sources

- Sparker: Sudden discharge of current between electrodes in water

generates seismic waves.

- Boomer: Current passes through Coil which moves a plate against water.

- Air Gun: High pressure bubble released in water.

- Aqua Pulse: Enclosed underwater chamber (elongated heavy-rubber

cylinder) filled with propane and oxygen. It is detonated by electric

spark. The explosion causes ballooning of the chamber which introduces

a pressure pulse in water.

6.4 Fold Coverage

We know that in reflection surveys the seismic waves hit the subsurface

depth point and are reflected back. During multiple shots with various shot

receiver combinations the same depth point may be hit multiple times and is

therefore referred as common depth point (CDP). Thus for a given spread of

geophones, the maximum number of times a CDP is hit by waves is called

fold coverage and is computed as follows;

Where C = Number of Channels

∆x=Picket Interval

∆g=Geophone Group Interval

∆s=Source Interval

6.5 Geophone Spread Geometries

The positioning of geophones along a 2D seismic profile with respect to the

source point is called the spread geometry. During acquisition the spread

moves one or multiple picket intervals along the seismic profile. Two

common spread geometries are discussed considering an 8 channel recorder.

1

2

x xFold C

g s

∆ ∆=

∆ ∆

Split Shooting

In this spread the shot is at the middle with equal number of geophones at

both sides. This spread remains symmetric throughout the acquisition profile

as shown below along with its stacking chart and fold coverage.

According to the figure the equidistance surface points where geophones or

source is placed are called pickets and the sub-surface points where the

waves hit are called CDPs. It can be seen that the CDP interval is half the

picket interval at the surface.

End-on Shooting

This spread starts in asymmetric form with the source at start of profile,

followed with half number of geophones at the forward side. As the spread

moves forward a geophone is added at the backward side. This is called roll-

in and it continues until the spread becomes symmetric. The spread

continues to move forward in symmetric form until the end of the profile is

reached. Finally at the end a geophone is removed from the forward side

with each forward step. This is called roll-out which continues until the

source reaches the end of the profile. The end-on shooting spread after roll-

out is shown, in the following figure along with its stacking chart.

A comparison of fold build-up, along a profile for both the above mentioned

shooting geometries is illustrated below. From the figure it is clear that end-

one shooting provides a better fold coverage along a profile.

Module 7

Seismic Noise

At the end of this module you would be

able to understand

� Difference between Signals and

Noise

� Coherent and Incoherent Noise

� Geophone Arrays

� Low and High Cut Filters

� Stacking to remove Incoherent

Noise

7.1 Signals and Noise

All events of interest are called signals while rest is called noise. Signals and

Noise are relative terms as in a certain set of analysis an event may be

considered as signal, while in another analysis it is considered as noise. In

seismic acquisition and processing our major emphasis is to enhance the

signals and suppress the noise. Thus our aim is to increase the Signal to

Noise Ratio (S/N). There can be multiple sources of noise. Noise may be due

to some other source or due to the same source responsible for the signals.

Seismic Noise is classified into the following two main types;

- Coherent Noise

- Incoherent Noise

7.2 Coherent Noise

It has a periodic pattern which can be followed at substantial distances along

the receiving profile. During seismic data acquisition the most common type

of coherent noise is Ground Roll.

Ground Roll These are surface waves primarily Rayleigh waves, having low-velocity and

low-frequency with relatively higher amplitude. They override the useful

reflections. In addition, refracted waves multi-reflected in a surface layer

and shear refractions are also encountered.

During acquisition ground roll is suppressed by two methods. As ground roll

has low frequency, usually below 10 Hz and our signals are above this

frequency, a low cut filter of 10-12 Hz is applied to suppress the ground roll.

In addition a group of geophones spaced at half the wavelength of the noise

and connected to a single channel also suppress ground roll as shown below.

Power Lines Induction If a seismic line crosses or passes close to a power line, then 50 or 60 Hz

coherent noise is induced into the nearby geophone channels. This noise can

be removed during recording by applying a notch filter.

7.3 Incoherent Noise

It has a random pattern and therefore also called Random Noise. There can

be multiple sources of random noise such as;

- Scattering from near-surface irregularities.

- Commonly occur when the shot point overlies or is close to gravels,

boulders or vuggy limestone all of which can cause scattering of waves.

- When stream banks and surface irregularities diffract energy.

Incoherent noise observed at one point on the surface is entirely unrelated to

that at another point. Similarly signals collected at the same point at different

times contain the same signal but different random noise. Thus, addition or

stacking of signals containing incoherent noise results in noise cancellation

as shown below.

7.4 Aliased Frequencies

If the analog signal contains frequencies higher than the Nyquist frequency

then such frequencies are aliased and appear as low frequencies after

digitization. Thus an anti-aliasing high cut filter of 125 Hz or above is

applied before analog to digital conversion.

+ + + =

7.5 Multiples

Sometimes seismic energy may be trapped by to and fro reflection between

interfaces. Thus a reflector appears twice in a seismic section, the second

time as a multiple at a greater time. The interference due to multiple

reflections appears similar to primary reflections and therefore sometimes

difficult to identify. Multiples are removed by predictive deconvolution. The

multiples have the same stacking velocity as the primary reflections, but

appear at a greater time. Thus they can be removed by avoiding them during

velocity picking.

There are several types of multiples as illustrated below:

Primary First

Order

Second

Order

Surface Multiples

First

Order

Second

Order

Interbed Multiples

Combination Multiples

Module 8

Seismic Data Processing I

At the end of this module you would be

able to understand

� Wave propagation through Earth

� Mechanical Processes

� Interactive Processes

� Gains, Spherical Divergence

� Band Pass Filter

� Deconvolution

8.1 Propagation of Seismic Waves through Earth

When Seismic waves propagate through Earth’s material they undergo

changes in their signature (waveform) due to the following phenomenon.

Geometric Spreading

Energy contained in a wave is proportional to the square of its amplitude. As

body waves propagate outward from a point source they spread spherically

with a constant energy. As the spherical wave front expands, the energy per

unit area must decrease as rapidly as the total area of the spherical surface

increases.

It follows Newton’s Inverse Square Law which states:

“The power per unit area in the direction of propagation, of a spherical wave

front varies inversely as the square of the distance from the source, assuming

there are no losses caused by absorption or scattering.”

Absorption & Attenuation: Convolution

There is also loss of amplitude due to absorption caused by frictional

dissipation of the elastic energy into heat. This loss from the source is found

to be exponential with distance as shown below. This absorption of energy is

called attenuation. Attenuation is directly proportional to frequency and

therefore higher frequencies are attenuated more than the low frequencies.

Thus low frequencies can penetrate further into the Earth.

Seismic energy may be absorbed due to several reasons some of which are

listed below:

- Crystal phase change

- Compaction of porous media

- Fracturing within the medium

- Friction due to relative motions of different parts of the same rock

- Viscous losses through fluid flow in a porous medium

- Temperature change through compression and dilation of the medium.

When a Seismic Wave propagates through Earth’s material it undergoes

convolution. Convolution is the term given to the mathematical technique

for determining a system output given an input signal and the system

impulse response.

Noise Addition

As a Seismic Waves from a desired source (Signal) propagate through the

Earth they may get mixed with waves from other undesired sources (Noise).

In addition undesired events from the desired source are also encountered.

Thus the recorded seismic traces undergo the following changes, which are

also illustrated below;

• Amplitude decay with distance, both horizontal (offset) as well as

vertical (depth)

• High frequencies absorbed

• Noise & unwanted events added

In seismic data processing, the main task is to enhance the amplitudes by

applying suitable gains, recover the lost high frequencies through

deconvolution and suppress different types of noise through band bass filter,

stacking and a number of other techniques. Thus our aim is to improve the

signal to noise ratio.

8.2 Mechanical Processes

Seismic data processing consists of several steps among which three

important steps called mechanical processes are always applied to the

seismic data. They are called mechanical processes as they change the

structure or form of seismic data, rather than performing some analysis to

change the amplitude, frequency or phase of the data. These processes along

with their input and resultant datasets are shown below.

Demultiplexing

We know that seismic recorder stores data in multiplexed form, where first

samples of all traces are stored first followed by second samples of all traces

and so on. Samples of all traces at a constant time make a time slice. Thus

the multiplexed data is in time slice order and we need to convert it into

trace sequential order, i.e. all samples of first trace followed by second trace

and so on. The demultiplexing of seismic data is illustrated below.

It can be seen from the previous figure that demultiplexing is simply

arranging the seismic data matrix from row (time slice) to column (trace

sequential) order. It must be noted that both multiplexed as well as

demultiplexed data are shot ordered and all traces belong to the same shot as

shown below. Modern recorders have field data units each with their own

analog to digital converter, instead of the conventional centralized recording

system. These systems store data in trace sequential order and thus there is

no need to apply demultiplexing to such data.

Sort

After performing some initial processing, the data needs to be sorted from

shot order to common depth point (CDP) order. In this process traces from

different shot receiver combinations that hit the same CDP are grouped

together thus forming a CDP ordered data as shown below. Again it can be

seen that this process is simply arrangement of data from one form to

another.

Stack

The number of traces in a CDP group depends on its fold coverage. For a 30

fold data the CDP order data will have 30 traces per group. After applying

dynamic corrections all traces in a CDP group are stacked (added) into a

single trace as shown below. This removes the random noise and enhances

the reflected events. Thus stacking changes the data form, but it is not purely

a mechanical process as it also improves the signal to noise ratio. It must be

noted that after stacking the data volume is considerably reduced. For a 30

fold dataset the data volume is reduced 30 times.

8.3 Interactive Processes

The seismic data processing workflow also includes some interactive

processes, which require a lot of human interaction. With the advancements

in graphics technology, computer aided interactive tools have been

developed to carryout these processes. The figure below shows the sequence

of these processes along with their usage application.

+ + + =

Trace Editing

This processing step is basically quality control of the input field data. The

input data may contain bad traces, reverse polarity traces and completely bad

and blowout records. Using an interactive environment, the data is viewed

record by record. All bad traces are muted (zeros), reverse polarity traces are

converted to normal polarity and bad records are killed. Thus all bad data is

removed before proceeding with further processing.

First Break Picking

In seismic refraction module, we discussed interactive and automated first

break picking for statics correction. The statics computed from field

refraction survey are referred as field statics. To further refine the data a

second set of statics corrections is applied to the data by interactively

picking refracted arrivals from all reflection records. This is referred as

refraction statics.

Velocity Picking

To stack the CDP ordered data dynamic corrections or normal moveout is

applied to the data. To apply dynamic corrections we need velocity

information for each reflector so that they are stacked properly. As velocity

changes laterally as well as vertically, velocity functions are interactively

picked at selected CDP groups. The picked velocity functions are

interpolated to apply dynamic corrections to all CDPs. More details about

velocity analysis procedure are given in the next module.

We now discuss some common processes which improve the recorded signal

quality.

8.4 Spherical Divergence Compensation and Gains

We know that as seismic wave front moves forward it experiences decay in

amplitude. This decay takes place in all directions, both laterally with offset

and vertically with depth as shown below.

In processing, the lost amplitudes must be recovered by applying some

spherical diversion compensation gain and other types of time variant and

trace balancing gains.

Time Variant Gain (TVG)

A time variant gain compensates the decay of amplitude with depth. Usually

a time variant function is computed and multiplied with the corresponding

samples of seismic trace.

It can be a linear gain function as shown below (left). The amplitudes of the

lower events have improved but they are still weaker than the top events.

The slope (m) and the intercept (C) can be adjusted to get better results.

Similarly we can have an exponential gain function shown below (right).

Here again the value of C must be optimized to get good results otherwise

for a large value of C we may get a bell-bottom type of trace with large gain

applied to deeper events.

TVG compensates the vertical decay of amplitude within each trace but it

does not take care of lateral decay of amplitude from near to far offset traces

as shown in the next figure. It can be seen in the figure that TVG has

enhanced the amplitudes of both near and far offset traces but the relative

difference in amplitudes of the near and far offset traces remains the same.

Trace Balancing

Certain applications, such as first break picking, require the maximum

amplitudes of all near to far offset traces to be balanced at a user defined

RMS amplitude (ARMS) level. This is achieved through trace balancing which

compensates for lateral decay of amplitudes. Thus trace balancing is a time

invariant type of gain. The procedure and application of trace balancing is

illustrated in the following figure.

The following figure shows the application of trace balancing to near and far

offset traces. It can be seen that the maximum amplitude of near and fat

offset traces in brought to the same level.

Automatic Gain Control (AGC)

The ultimate gain is a combination of TVG and trace balancing. Instead of

scanning the complete trace to get AMAX , the trace is scanned within a sliding

time window (operator length). Thus several gain factors (GF) are computed

at the center of each window and joining them gives a time variant gain

function as shown. It is called automatic gain control as no coefficient needs

to set and gain factors are automatically set according to the amplitudes in a

time window.

8.5 Band Pass Filter

Seismic data is band limited which may range from 10 - 125 Hz. Usually

the dominant frequency is 35 - 45 Hz, thus a band pass filter (BPF) of 10-80

Hz can be applied to the data to pass all useful signal frequencies and

suppress the remaining frequencies.

A BPF can be considered as a combination of low-cut and high-cut filters to

pass a band of frequencies. Thus a BPF allows all frequencies between the

low and high cut off frequencies as shown below.

The working of BPF is given in the following block diagram. A filter

operator wavelet is generated according to the low-cut (fL) and high-cut

frequencies (fH). The resulting wavelet contains all frequencies within these

limits. This wavelet is convolved with the input trace to get the output

filtered trace.

8.6 Deconvolution

When seismic signals convolve with the Earth’s material, high frequencies

are absorbed. These lost high frequencies are necessary for improving the

temporal resolution, thus they must be recovered. Deconvolution is a

mathematical process used to reverse the effects of convolution on recorded

data. It is exactly what it sounds like: the undoing of undesired convolution.

During convolution high frequencies are attenuated. In deconvolution we try

to re-introduce these lost frequencies. Thus deconvolution can be considered

as an inverse filter. Some common types of deconvolution are spiking

deconvolution, predictive deconvolution and surface consistent

deconvolution.

The working of deconvolution is given in the following bock diagram. The

input trace is converted into frequency domain using Fast Fourier transform.

The inverse of amplitude spectrum is computed. Then pre-whitening is

added to the inverse spectrum to avoid zeros. This is done by adding one to

the amplitude of all frequencies in the spectrum. The phase spectrum is set

to zero, this makes the resultant wavelet operator as zero phase. The final

amplitude and phase spectrum are transformed back to time domain by using

Inverse Fourier transform to get the inverse wavelet operator. Finally this

operator is convolved with the input trace to get the output deconvolved

trace.

Module 9

Seismic Data Processing II

At the end of this module you would be

able to understand

� Velocity Analysis / Constant

Velocity Stack

� Dynamic Corrections / Normal

Moveout

� Stacking: Raw, Brute Stack

� Residual Statics and Final Stack

� Migration

9.1 Seismic Data Processing Flow

In the preceding module we discussed some mechanical, interactive and

basic processing operations. In this module we will focus on some other

important processing functions and consider the complete seismic data

processing sequence. A generalized seismic data processing flow is given in

the following block diagram. It also shows the mechanical, interactive and

basic processing functions already discussed.

The field data is demultiplexed followed by geometry setup, where spread

layout, navigation data and field statics are updated to seismic data headers.

Then trace editing is performed and all basic processing such as geometric

spreading compensation, filters and deconvolution are applied before sorting

the data to CDP order. A copy of the geometry applied data is sub-sampled

to 8 milliseconds for first break picking. Refraction statics are computed and

can be applied to pre or post sorted data. Then normal moveout correction

(NMO) is applied by using a regional velocity function and the data is

stacked to get a raw stack. In the next stage velocity analysis comprising of

constant velocity stack (CVS) and velocity picking at selected CDP

locations is performed and the picked velocity functions are used in the

NMO to get a Brute stack. In the final go the velocities may be further

revised and residual statics are applied to get the final stack. The structures

in the final stack are migrated to their true positions to get the migrated

stack.

9.2 Dynamic Corrections

Dynamic corrections or Normal Moveout (NMO) are corrections applied to

CDP ordered data to reduce all source receiver slant travel times (Tx) into

zero offset vertical times (To) as shown below.

It can be seen in the figure, as the offset increases Tx also increases. Thus the

CDP family traces represent a hyperbolic travel time curve for a reflector.

As we need to stack these traces the slant travel times must be aligned along

a straight line by reducing them to vertical times. Thus the NMO correction

is simply reducing the data to zero offset, given by;

In the above figure Tx, To and offset x, make up a triangle. Now the problem

is that two sides of the triangle have units of time while the third side, the

offset, has units of distance. To solve this triangle we need to convert x into

time. For this we need velocity of the reflector, called NMO velocity (VNMO),

which is obtained from velocity analysis. Thus from the Pythagorean

Theorem we have;

NMO x oT T T∆ = −

2

2

2 2

NMO

xx o V

T T= +

Now the NMO correction becomes;

After applying these corrections, all CDP traces are aligned and can be

stacked. Like static corrections, the dynamic corrections are also in the form

of a time shift. In static corrections a constant time shift is applied to all

samples of a trace, thus the whole trace is moved up or down along time

axis, but the relative time gap between events remains static. On the other

hand, dynamic corrections are computed and applied to each reflector. As

the velocity of each reflector varies, its NMO correction also varies and

therefore each reflector is moved at a different rate. Thus there is a

stretching or shrinking of time gap between events, therefore these

corrections are called dynamic corrections.

Let’s consider a reflector with NMO velocity 2250 M/Sc. If the appropriate

NMO velocity is used the events are aligned, if a lower velocity is used the

events in CDP gathers (traces) are stretched up, called over-corrected and if

a higher velocity is used the events remain under corrected, as shown below.

9.3 Velocity Analysis

From the preceding section, we know that NMO corrections require velocity

information. In seismic data processing velocity analysis is an important step

in which velocities are picked from seismic data at selected CDP locations.

The CDP should not be selected at uniform intervals. Suitable locations are

2

2

2

NMO

xNMO o oV

T T T∆ = + −

the crests and troughs of a folded area. Zones of poor signal to noise ratio,

faults and near-surface anomalies must be avoided for velocity analysis as

shown below.

Constant Velocity Stack (CVS)

In real Earth, the velocity generally increases with depth, thus each reflector

will have a different velocity. In velocity analysis out task is to determine

the appropriate velocity of each reflector. As velocity also changes laterally,

we need to select several CDP locations where velocity analysis is to be

performed. To view the continuity of reflectors we need a group of CDPs,

thus the number of CDPs on both sides of the selected CDP location is

specified. If 10 CDPs are specified on both sides, the selected location will

have a group of 21 CDPs. Finally we also need to specify the minimum and

maximum velocity range we expect in the region, and a velocity increment.

Typical values can be 1500-5000 M/Sec with an increment of 100 M/Sec.

The CVS method uses a constant velocity in NMO and stacks all CDP

traces. The constant velocity is iterated from minimum to maximum range

with the specified increment. In this way we get velocity panels for each

selected CDP location, showing CDP group traces stacked with a range of

velocities as shown in the next figure. From NMO we know that each

reflector will stack with strong amplitude if its appropriate velocity is used.

Higher or lower velocities will not stack the reflector properly. As velocity

increases with depth the shallow reflectors will stack well at lower velocities

while the deeper reflectors at increasing velocities. Thus on a velocity panel

we mark a point for each reflector where it is best stacked. These points are

velocity-time pairs and joining them gives us a velocity function which

increases with time.

In this way we pick velocity functions for each selected CDP location. The

picked velocity functions are interpolated for in between CDPs and used in

NMO corrections to get a Brute stack. These velocities are referred as root

mean square (RMS), NMO or Stacking velocities, as all three types have

approximately same values. It must be noted that NMO/Stacking velocities

can be successful if they vary within ±20% of their true value.

9.4 Residual Statics

If our main horizon of interest does not show a good continuity in the Brute

stack, we may need to apply residual statics. In this technique we need to

mark a point on the horizon and specify its minimum and maximum times in

the section. It correlates the horizon events in all traces and generates a

smooth trend of the horizon. It applies a plus-minus shift to all traces, so that

the horizon events are aligned according to the smooth trend. Now the

horizon shows a continuous trend. It must be noted that by using this

technique our horizon of interest becomes continuous while other horizons

may get disturbed.

9.5 Migration

The final seismic section does not represent the true geometry of the

reflectors, thus migration is applied to the final stack with the following

main objectives.

Move Dipping reflectors to their true position

Corrects for the geometric displacement of data from a dipping reflector

and/or lateral velocity changes and places them in their true spatial position

rather than at an assumed point in depth between the source and the receiver.

Collapse Diffraction Patterns

Eliminate the signal interference caused by point diffractors.

To understand why seismic expression does not show the true position of

events, consider the following figures.

At the surface the common mid point (CMP) lies at the middle of shot and

receiver positions. For a horizontal bed the common depth point (CDP) lies

exactly below the CMP.

Now for a dipping bed the CDP does not lie below the CMP as shown.

The seismic section will still show the CDP below the CMP, thus we need to

shift it to its true position.

Thus an anticline will appear larger with wider flanks in seismic as shown.

Similarly a syncline will appear as a bow-tie in seismic as shown.

Migration is an extremely compute intensive process. Several migration

algorithms have been developed, some common types are listed below:

- Kirchhoff’s Migration

- Finite Difference (FD) Migration

- Frequency Wave number (FK) Migration

- Reverse Time Migration

- Stolt Migration

With the availability of large computing power, migration can also be

applied to pre-stack data in one of the following forms:

- Pre Stack Time Migration (PSTM)

- Pre Stack Depth Migration (PSDM)

Module 10

Seismic Resolution

At the end of this module you would be

able to understand

� Temporal Resolution: Frequency

and Bandwidth

� Spatial Resolution: Picket Interval

and Fresnel Zone

� Phase Uncertainty

� Signal to Noise Ratio

10.1 Resolution

The smallest object that can be imaged or resolved by a system or technique

is called the resolution of the system. Such as telescopes and microscopes

have resolving power which decides the smallest object that can be view by

these instruments.

In Earth sciences resolution implies to the vertical thickness and lateral

extent of a subsurface geological body that can be delineated with a

geophysical method.

10.2 Seismic Resolution

Seismic resolution refers to the ability of the seismic method to image the

thinnest and smallest subsurface objects. Seismic resolution is of two types;

Temporal Resolution

It is the vertical resolution which accounts for the thickness of sub-surface

beds that can be resolved. It depends on frequency of seismic waves and is

1/4th of the wavelength. Thus higher the frequency, the smaller will be the

wavelength and ultimately we get a higher resolution as shown below.

To understand seismic resolution a three layer model, consisting of a thin

layer sandwiched between two thick layers is given in the next figure. It can

be seen that with low frequency the thin bed is not resolved, but using a

higher frequency the bed is clearly resolved.

Spatial Resolution

It is the lateral resolution which accounts for the lateral extent of subsurface

bodies that can be resolved. It depends on spacing of sensors (geophones) on

the surface. We know the CDP interval is half the picket interval. Thus the

sub-surface target object must be multiple times larger then the CDP

interval, if it is smaller than the CDP interval it will be missed out. Consider

the following example where a small lens shaped object is shown. For large

sensor spacing only one ray path hits the body and therefore its shape cannot

be delineated. If the sensor spacing is reduces the CDP interval is also

reduced and several ray paths hit the body at multiple CDPs thus its shape is

clearly identified.

10.3 Fresnel Zone

Horizontal resolution is defined in terms of Fresnel Zone which indicates

how close two adjacent points in the subsurface can be, while still being

distinguished from one another. We know that vertical resolution is 1/4th of

the wavelength of frequency. The horizontal resolution also depends on

wavelength of frequency and depth of the target. The Fresnel zone and its

relationship with horizontal resolution is illustrated in the next figure.

The figure shows two wave fronts separated apart by 1/4th wavelength of the

dominant frequency. The upper wave front is at a depth Z. Now AA/ is the

Fresnel zone and its half is the Fresnel Zone radius, which is the horizontal

resolution. Mathematically the Fresnel Zone Radius is given by;

It can be seen that with the increase in depth the wave front expands and the

Fresnel zone radius also increases, thus the resolution decreases.

10.4 More on Seismic Resolution

Resolution is our ability to clearly interpret the nature of reflectors from a

limited seismic display. The resolving power of seismic data is limited by

the following:

Bandwidth

Previously we discussed that vertical resolution depends on frequency, but in

actual it depends on bandwidth, which is a range of usable frequencies

contained in seismic data. Bandwidth is not simply the difference in

frequency from high to low limits. It is the logarithm of the ratio of the

frequency limits given by;

For base 10 we have;

/ 2F

R Zλ=

2 ( )h

l

fBandwidth Log

f=

10

10

( )

(2)

h

l

fLog

fBandwidth

Log=

The unit of bandwidth is Octave, which simply represents doubling of

frequency range.

10 Hz – 20 Hz is one Octave

20 Hz – 40 Hz is second Octave

40 Hz – 80 Hz is third Octave

Thus a signal containing all frequencies from 10 Hz to 80 Hz has a

bandwidth of 3 Octaves.

The figure below shows the effect of bandwidth on resolution. It can be seen

that with the increase of upper frequency from 60 to 140 Hz the bandwidth

increases and thus the resolution increases, but on increasing the lower

frequency from 8 to 120 Hz the bandwidth decreases and ultimately the

resolution decreases, in spite of the fact that high frequencies are present.

Thus it not only the high frequency, but a range of low to high frequency

called bandwidth, which improves the resolution. A minimum 2.5 Octaves is

required for a good seismic resolution.

Phase Uncertainty

Zero phase data provides the simplest expression of reflection events as side-

lobe interactions are minimized. For interpretation purposes the data should

be reduced to zero phase. It can be seen in the following figure that zero

phase data shows the exact position of the reflector, while in non zero phase

data we are not sure about the precise position of the reflector.

Signal to Noise Ratio (S/N)

Random Noise can seriously interfere with the resolving power of data.

Strong Noise can inhibit the ability to see major reflections. Thus for a better

resolution, the data must have a good S/N, as shown below. When S/N

decreases, the signals are masked by the noise. A S/N=1 implies that both

signal and noise amplitudes are same

Module 11

Seismic Interpretation

At the end of this module you would be able

to understand

� Seismic Section: Display Modes, Vertical

and Horizontal Scales

� Marking Horizons and Faults

� Base Map and Contouring

� Seismic Velocities and Time to Depth

Conversion

� Seismic Modeling

� Sonic Log, Density Log and Synthetic

Seismogram

11.1 Seismic Data Display Standards

Seismic data can be displayed in a number of industry standard formats. A

sample of each display type is given below in the form of an individual trace

and a complete section.

Wiggle + Variable Area: Waveform with shaded crests

Wiggle: Only Waveform.

Variable Area: Only shaded crests

Color Attributes: Waveform with amplitude based colored crests

Colored Density: Crests in Red, Troughs in Blue (or other colors)

11.2 Seismic Section Display Scales

Seismic sections are displayed with separate horizontal and vertical scales.

Setting up of these scales is critical from interpretation point of view as

sometimes gentle dips may appear as horizontal beds by using a compressed

vertical scale. Similarly the dip may be exaggerated by using a large vertical

scale. Setting up of both these scales is discussed below along with

examples.

Horizontal Scale

The horizontal sale is described in the form of Traces per Inch (TPI).

Increasing the number of traces per inch reduces the seismic section

horizontal scale as more trace are packed within one inch as shown below.

Vertical Scale

The vertical scale is described in the form of Inches per Second (IPS). In this

case increasing inches per second enlarges the seismic section vertical scale

as shown in the next figure.

Typical setting for seismic display scales can be 24 TPI and 2.5 IPS.

11.3 Base Map

In addition to seismic sections, base map is also an important component of

interpretation, as it displays the spatial position of each picket of a seismic

section. Its also shows the spatial relationship of all seismic sections under

consideration, their tie point locations and provides the framework for

contouring. Seismic base maps have also been standardized as shown below.

The line is annotated on both sides in the direction of line, while pickets are

annotated perpendicular to the line. Picket interval and picket annotation

interval are specified before generating a map.

The base map is produced using some projection system such as Lambert

Conic projection or Universal Transverse Mercator (UTM) projection. The

base map also shows grids of geographic latitude-longitudes and/or

projected grid coordinates. The base map is produced at a specified scale,

such as 1:10000, which represents the number of real Earth units that equal 1

unit on the map.

11.4 Seismic Interpretation

In interpretation our main task is to identify various reflectors or horizons as

interfaces between geological formations. For this good structural and

stratigraphic knowledge of the area is required. Thus during interpretation

we mark the horizons and faults on the seismic section. Initially

interpretation was done manually on paper sections, but with the availability

of powerful computer systems with graphics support, computer aided

interpretation systems are being used in the industry as shown below.

Like first break picking the horizons can also be marked in the following

four modes, but the selected mode must be used throughout the project.

The complete seismic interpretation workflow is given in the following

figure. Accordingly all seismic sections in the projection are interpreted by

marking horizons of interest and faults. The marked horizon and fault times

along with their CDP numbers and X,Y navigation coordinates for all

interpreted sections are output to the gridding and contour module which

generates a contour map or 3D surface of the horizon. If there is a

prospective zone, a well point is marked on the respective section and the

contour map.

A time contour map of a horizon, along with faults and seismic lines is given

in the next figure.

Using seismic velocities this time contour map is also converted into depth

map.

11.5 Time to Depth Conversion

The interpreted seismic section is in time domain. In order to get a true

geological picture it must be converted into a depth section. This conversion

requires reliable velocity information which varies vertically as well as

horizontally. The main source of velocity information is seismic velocity

picked during data processing. More accurate velocities can be obtained

from check shots and vertical seismic profiling (VSP) surveys. Using Dix

equations, the RMS velocities, from seismic, are converted into interval and

finally average velocities. These velocities are used for time to depth

conversion as shown below. The horizon times are two way times (TWT)

therefore they are divided by two before depth conversion.

It must be noted that the vertical axis, in the above figure, is now depth in

meters. The Dix equations for conversions between RMS, interval and

average velocities are given below.

RMS to Interval Velocity

2 2

1 1

1

int i i i i

i

i i

Vrms T Vrms TV

T T

− −

−

−=

−

Interval to RMS Velocity

Average to Interval Velocity

Interval to Average Velocity

11.6 2D Seismic Modeling

Sometimes we have a geological cross-section and we want to know the

seismic response of this section, which may be used for deciding parameters

in planning a new seismic survey. Similarly, we have interpreted a seismic

section in which the horizons and faults make up a geological section and we

want to move back and generate its seismic section. This is done to confirm

our interpretation. The generation of a seismic section from a geological

section is called 2D modeling. It is the reverse of seismic interpretation as

shown below.

2

1

1

int ( )n

i i i

i

i

i

V T T

VrmsT

−=

−

=∑

1 1

1

int i i i i

i

i i

Vave T Vave TV

T T

− −

−

−=

−

1

1

int ( )n

i i i

ii

i

V T T

VaveT

−=

−

=∑

The modeling procedure involves digital signal processing techniques. We

generate a source wavelet which mathematically represents our real sources

like dynamite or vibroseis. There are several techniques for generating

wavelets such as Ricker wavelet, Klauder wavelet and Summed wavelet.

The figure below shows a software interface to setup parameters for

generating a wavelet.

To model a geological cross-section it must be in some digital format. We

assign reflection coefficients to various horizons in the cross-section on the

basis of their velocity and density contrasts. The acoustic impedance (I) of a

layer is given by;

I V ρ=

where V is the velocity and ρ is density.

Now the reflection coefficient (RC) is given by;

1 1

1 1

i i i i

i i i i

V VRC

V V

ρ ρ

ρ ρ− −

− −

−=

+

After assigning the reflection coefficients to all horizons the section it is

convolved with the source wavelet by specifying a CDP interval. A synthetic

seismic section is generated as a result of this modeling process as shown in

the next figure.

11.7 Synthetic Seismogram

In the previous section we discussed 2D modeling, now synthetic

seismogram is basically 1D modeling. In this procedure we also generate a

source wavelet. Now instead of a 2D geological cross-section we have

petrophysical logs; Sonic (DT) and Bulk Density (RHOB) logs which

respectively provide the velocity and density information of subsurface

layers. The DT is a delay time log and its inverse gives the velocity. These

logs are acquired in the borehole. We use this velocity and density data to

compute a series of reflection coefficients called reflectivity series. This

series is convolved with the source wavelet to get a synthetic seismogram. In

this case we have performed the convolution with only one reflectivity series

(1D), thus only one seismic trace is generated. Graphically we plot multiple

copies this synthetic trace so that it appears like a stack section as shown in

the next figure. The synthetic seismogram vertical units are meters or feet

and it can be converted into time units by using its own velocity information.

Synthetic seismogram is matched with the seismic section at the well point

to correlate the succession of reflectors. It may also be used to calibrate our

seismic velocities.