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Intermediate Training in Meteorological Instrumentation

and Information System

Lecture Note on Seismology: Phase-I

Self Study / E Learning

Centre for Seismology (For use in IMD)

Note: Some of the figures are taken from the book ‘An Introduction to Seismology, earthquakes

and earth structure by Seth Stein and Michael Wysession’ and ‘http://earthquake.usgs.gov/’

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Syllabus

Introduction to Seismology –

Internal structure of the Earth, Plate tectonics, Physics of earthquake processes; Types of faults

and fault mechanisms; Seismicity and Seismotectonic features of India. Elastic Wave theory:

Seismic wave propagation & characteristics, Travel-time tables and Velocity models. Earthquake

source parameters; Magnitude, intensity, energy; etc.; Earthquake statistics; digital data analysis

and location of earthquakes; Seismological operations and information dissemination.

*****************

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1. Introduction to Seismology:

Seismology is the science of Earthquakes (derived from Greek word Seismos meaning

Earthquake and Logos meaning science) and related phenomena. Earthquakes are the

manifestations of sudden release of energy accumulated over extensive periods of time in the

upper part of the Earth. The energy thus released is radiated as seismic waves of various types,

which propagate in all directions through the Earth’s interior and are recorded by sensitive

instruments, called seismographs, installed on or near the Earth’s surface. The seismogram thus

produced reflects the combined effects of the source, the path, the characteristics of the recording

instrument and the ambient noise at the recording site. Through the study of seismograms, it is

thus possible to extract information pertaining to the internal structure of the earth, the earthquake

source characteristics and the ambient noise characteristics of various sites. It is worth mentioning

here that the overwhelming part of contemporary knowledge on the internal structure of the Earth

and its dynamics has been inferred from seismological evidence.

Unlike the science of Meteorology where weather forecasting is possible, the science of

Seismology has not yet reached a stage, where reliable prediction of earthquakes is possible in terms of space, time and magnitude. However, efforts are continuing in India and abroad towards

earthquake prediction research and monitoring of earthquake precursors, which could prove to be a promising tool in developing a forecasting technique in future.

While efforts towards earthquake prediction should continue, it is equally important to

take necessary steps to minimize the damage caused to property and lives by these natural

calamities. Based on the past history of earthquakes combined with other geological information,

Bureau of Indian Standards (BIS) has classified the country into four seismic zones (Zones-II, -

III, -IV and -V). This Seismic Zoning map provides broad guidelines towards assessment of

expected seismic intensities in a given region. BIS has also published criterion for construction of

earthquake resistant structures, which if strictly followed will ensure substantial reduction in

damage to structures and loss of life in the event of an earthquake. Further, losses due to

earthquakes can also be considerably reduced through proper planning and implementation of pre-

and post- disaster preparedness and management strategies. Timely dissemination of information,

to concerned state and central government agencies, pertaining to occurrence of earthquakes also

plays a very important role in dealing with the post-disaster relief and rehabilitation related

matters.

1.1 Internal structure of the Earth:

The interior of the Earth is mainly divided into three layers called Crust, Mantle and Core. The

average radius of the Earth is 6370 km.

The uppermost part of the Earth, called Crust extends down to a depth varying between 5 and 70

kms. and the boundary separating the Crust and the Mantle below is called as Mohorovicic

discontinuity (abbreviated as Moho). The depth of the Moho varies between 5-8 km. in oceans,

25-40 km. below continents and 60-70 km. under mountains. The Crust is again divided into

Granitic (Top) and Basaltic (Bottom) layers based on the composition of the material. The boundary between the Granitic and Basaltic layers is called the Conrad discontinuity. In studies

of nearby earthquakes (spatial distance 1000 km.), we often assume a crust consisting of two horizontal layers of approximately the same thickness, separated by Conrad. For a typical crustal

model under the deep ocean, we normally omit the granitic layer. The Crust is mainly made of material rich in Silica, Alluminium and Magnesium. There is a sudden increase in the velocities

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of Compressional waves from 6.5 km./s at the base of the Crust to about 8.0 km./s in the

underlying mantle rocks.

Figure 1.1: Internal Structure of Earth

The layer below the Crust is called Mantle and extends down to the Core boundary. The whole of the Mantle is now considered to be essentially solid and to a large extent radially

homogeneous. The Mantle is also divided into two distinct layers called Upper Mantle and Lower Mantle depending on the temperature, density and chemical composition of the material. The

Upper Mantle extends from Moho to a depth of about 700 km., where the compressional wave velocity gradient suddenly decreases. The Lower Mantle below extends to the Mantle-Core

Boundary at a depth of about 2900 km. The compressional wave velocity increases from about

8.0 km./s just beneath the Moho to 13.7 Km./s at the Core-Mantle boundary. One of the important

features of the upper Mantle is the world-wide existence of a Low Velocity Layer (LVL)

between about 100 and 250 km. below the surface. Within the LVL, the rocks are partially

molten, the rigidity is low, the attenuation is the largest and the seismic wave velocities fall by

about 6%. There is seismological evidence for a few other interfaces also at depths of 400 kms.

and 650 kms, although they are not as sharp as other discontinuities. The Mantle material is rich in

Olivine.

The third and the innermost layer of the Earth is called Core. The Core is also divided into

two parts as Outer Core and Inner Core. The Outer Core extends from the Mantle-Core

boundary to a depth of about 4980 km. The layer below the Outer Core is called the Inner Core.

The Core-Mantle boundary is again marked by a steep fall of the compressional wave velocity

from 13.7 to 8.1 km./s and cessation of shear waves. It is known that shear waves cease to exist at

this depth due to the fluid character (no resistance to shear, i.e. no shear strength) of the Core. The compressional wave velocities in the Inner Core are significantly higher than that in the outer

Core and shear waves are transmitted through the Inner Core. The Core is made up of Nickel and

Iron.

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1.2 Plate tectonics:

All places on the Earth are not equally seismic. Earthquakes are generally found to occur

along specific regions called “Seismic Belts”. There are three main belts around the globe along

which majority of earthquakes have occurred. They are: a) the one which runs around the Pacific

ocean and includes the western coasts of south America & north America, the eastern coast of

Asia, islands of southeast Pacific and New Zealand. This belt is called “Circum Pacific Belt”, or

“Ring of Fire”, due to the reason that majority of the volcanoes of the world also occur along this

belt, b) the belt that runs from south Pacific islands through Java, Sumatra, central Asian

mountains, Caucasus mountains, Greece, Italy and Spain and c) the one that runs from north to

south in the middle of the Atlantic ocean. In India, the main seismic zone runs along Himalayan

mountain range, northeast India, Andaman-Nicobar islands and Rann of Kutch region.

Figure 1.2: Plates Tectonic

According to the theory of Plate Tectonics, the uppermost part of the Earth is considered

to be divided into two layers of distinct deformational properties. The upper rigid layer, called the Lithosphere, consists of crust and a part of rigid upper Mantle material. The thickness of this

layer varies between 50-100 km, depending upon whether the crust is made up of oceanic or continental material. The lower layer, called the Asthenosphere, consists of less brittle (more

deformable) Mantle material and extends down to a depth of about 700 km. The rigid lithospheric shell is broken into, say, about a dozen or so irregularly shaped pieces, called major Plates (not

coinciding with continents) and a large number of minor or secondary Plates. These lithospheric

plates are not stationary; but are in constant motion with a velocity varying between 2-10

cm./year. In other words, the rigid lithospheric plates float on the less rigid, soft asthenosphere

underneath in a complex manner. Convection in the Earth’s mantle driven by the heat generated

from within the Core is assumed to be the driving force for the movement of the lithospheric

plates over such large distances.

The theory of plate tectonics was originally proposed in 1912 by a German scientist,

A.Wegner and is based on the following hypotheses:

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• Continental drift – Physical matching of continental boundaries of eastern coast of south America and western coast of Africa.

• Sea-floor spreading, Geomagnetic reversal and Polar wandering.

• Palaeo-Climatic studies of various continents. The edges of the oceanic and/ or continental plate boundaries mark the regions of

destructive earthquake activity and volcanic activity. The plate boundaries may be classified into

one of the three following categories, depending upon the type and nature of movement across

these boundaries:

a) Constructive Plate boundaries, are those where two plates move away from each other

due to forces of magma across the ridge axis. The opening where two plates diverge is continuously filled by ascending mantle material, as in the case of Mid-Oceanic ridges.

b) Destructive Plate boundaries, are characterized by converging plates. As two plates

converge, one plate usually bends beneath the other and descends into the soft, hot asthenosphere, a process often referred to as subduction. The descending slab, also called as subduction zone or

Wadati-Benioff zone, assimilates with the surrounding mantle at a depth of about 700 km, i.e. approximately at the lower limit of the asthenosphere, due to temperatures and stresses existing at

that depth. Oceanic trenches are the best example of subduction zones. When two continental

plates collide with each other, one plate will thrust below the other they give rise to a collision

zone, as is the case with Indo-Eurasian Plate boundary.

c) The third type of plate boundaries link the above referred mobile belts into an

interconnecting network. These are the boundaries where the plates are neither created nor

destroyed, but slide past each other horizontally. They may be seen as the interconnecting

Transform faults in the Mid-oceanic ridge regions.

Figure 1.3: various types of plate boundaries

1.3 Physics of earthquake processes;

Earthquakes are usually caused when the underground rocks suddenly break along a plane

of weakness, called fault. A fault is nothing but a crack inside the Earth.

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Figure 1.3.1: Geographic fault

Prof H.Reid proposed the elastic rebound theory of earthquakes on a fault. In this model,

materials at distances on opposite sides of the fault move relative to each other, but some friction

on the fault “locks” it and prevents the sides from slipping. As the tectonic forces continue to

prevail, the plate margins exhibit deformation as seen in terms of bending, compression, tension

and friction. Eventually the strain accumulated in the rock is more than the rock on the fault can

withstand, and the fault slips resulting an earthquake. The build up of stresses and subsequent

release of the strain energy in the form of earthquakes is a continuous process, which keeps on

repeating in geological time scale.

Figure 1.4: elastic rebound model of earthquake process.

Following are some theoretical assumptions that explain the forces, which cause

accumulation of stresses inside the earth:

• Drifting of continents and mountain building process.

• Shortening of Earth’s crust due to cooling and contraction.

• Disturbance of mass distribution on the Earth’s surface as a result of erosion of high lands and deposition of sediment in the sea.

• Generation of heat by radioactive material inside the Earth’s crust.

1.4 Types of faults and fault mechanisms;

Depending upon the type and direction of movement of blocks of rock, faults may be classified as:

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• Normal faults, are the faults caused due to sliding of one block downward and away from another block of rock. These faults usually occur in areas where a plate is very slowly

splitting apart or where two plats are pulling away from each other. Earthquakes occurring in Mid-oceanic ridges show normal fault mechanism.

• Reverse faults, are the faults produced when one plate pushes into another plate. They also

occur where a plate is folding up because it is being compressed by another plate pushing

it. At these faults, one block of rock slides underneath another block or one block pushes

up over the other. Earthquakes occurring in Himalayan region are characteristic of reverse

fault mechanism.

• Strike-slip faults, are produced when two plates slide past each other. The blocks of rock

along these faults move in a horizontal direction. The best example of strike-slip fault is the San Andreas fault in California, which produced several destructive earthquakes in the

past. To describe the fault geometry, we assume the fault is planar surface across which the

relative motion occurred during an earthquake. In this model the fault plan is characterized by its normal vector. The motion is given by the slip vector in the fault plane. The slip vector indicates

the direction in which the upper side of the fault (hanging wall block) moved with respect to the lower side ( foot wall block). The direction of the motion is represented by the slip angle λ,

measured counter clockwise in the fault plane from x1 direction, which gives the motion of the

hanging wall block w.r.t. the foot wall. The orientation the system relative to that geographic one,

the fault strike Фf is defined as the angle in the plane of the earth's surface measured clockwise

from north to the x1 axis.

Figure 1.5: Fault Geometry. Although slip direction varies such that the slip angle varies form 0 to 3600, several basic fault

geometry described by special values of slip angle.

• When two sides of the fault slide horizontally by each other, pure strike-slip motion

occurs.

– When λ=00 the motion is called left –lateral

– When λ=1800 the motion is called right–lateral

Figure 1.6: Pure strike-slip faults.

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• Other basic motion is dip slip.

– When λ=2700 the motion is called normal faulting.

– When λ= 900 the motion is called reverse or thrust faulting.

Figure 1.7: Pure dip-slip faults.

1.5 Seismicity and Seismotectonic features of India.

All places on the Earth are not equally seismic. Earthquakes are generally found to occur

along specific regions called “Seismic Belts”.

Figure 1.8: Global seismicity map clearly delineating tectonic plates

There are three main belts around the globe along which majority of earthquakes have

occurred. They are:

• Circum Pacific Belt, or Ring of Fire,

• Alpide belt and

• north to south ridge in the middle of the Atlantic ocean. Seismicity of India is mainly due to the Himalayan seiimic belt wich is in fact a part of the

world-wide Alpide belt extending from Indonesian islands to Spain through Burma, Himalayas, middle-east and southern Europe. The great Himalayan mountain range extending from Kashmir

in the west to the hills of Assam and upper Burma in the east is a region, which has witnessed

several most destructive earthquakes of the world, which were responsible for the loss of life and

property. The geology of the region is highly complex in which the younger foothills comprising

of the Siwalik series have been thrust over by the old Himalayan rocks along what is known as the

great Himalayan Boundary Fault. The thrusting has been attributed to the southward pressure of

the Asian plate against the Indian peninsula. According to the theory of Plate tectonics, the large

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Himalayan mountains are attributed to the continent-continent collision, where the single trench zone of plate consumption is replaced by a wide zone of deformation due to cracking and

splintering of the lithosphere and are characterized by a single dominant direction of thrusting. The western boundary of the Indian plate is marked by the Carlsberg Ridge, the Owen fracture

zone and the Kirthar-Sulaiman shear zone in west Pakistan. The differential movement between

the plates produces the observed left-lateral displacement along the Kirthar-Sulaiman shear zones,

if the Indian plate is assumed to be moving northeastward faster than the Afghanistan plate. The

eastern boundary of the Indian plate passes through Assam-Burma border to Andaman-Nicobar

islands. Focal mechanism solutions of earthquakes have shown thrust faulting along the Main

Boundary Fault. Strike slip and normal faulting is also indicated showing complexity of the

regional tectonics.

The Rann of Kutch and Cambay region is also comparable in seismicity with Himalayan

region. A number of faults aligned mainly east-west are located in the region and have contributed

to it’s seismicity. The faults apparently terminate against the western side of Cambay trough,

which in turn runs north-south and is bounded by faults on the eastern and western sides. Various

parts of Gujarat region have in the recent past witnessed swarm type activity and or explosion like

sounds. The Indian Peninsula is a Precambrian shield consisting of Archean rocks, which are mostly covered by recent sediments like Cuddapah, Vindhyans, etc. The central portion of the

Peninsula experienced the greatest uplift. The Peninsular Shield region is known for its intraplate seismicity, as demonstrated by the recent earthquakes viz., Latur earthquake of 1993 and Jabalpur

earthquake of 1997. The Bhuj earthquake of 26th

January, 2001 is another significant intraplate event in the modern instrumental era.

Figure 1.9: Seismicity map of India and neighborhood region with magnitude 5 and above

(upto June, 2011) *****************

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2 Elastic Wave theory:

Seismic waves are elastic waves. Earth material behaves elastically to transmit them. The degree of elasticity determines how well they are transmitted. In an underground explosion or in

an earthquake rupture, the surrounding Earth material is subjected to stress (compression, tension

and/or shearing) and a consequence, it undergoes strain, i.e., it changes volume and/or distorts

shape. In an inelastic (plastic, ductile) material this deformation remains while in elastic media,

the material returns to its original volume and shape when the stress load is over.

The degree of elasticity/plasticity of real Earth material depends mainly on the strain rate,

i.e., on the length of time it takes to achieve a certain amount of distortion. At very low strain

rates, such as movements in the order of mm or cm/year, it may behave ductile material for

example formation of geologic folds or the slow plastic convective currents of the hot material in

the Earth’s mantle with velocity on the order of several cm per year. However, Earth reacts

elastically to the small but rapid deformations caused by a transient seismic source pulse. Within

its elastic range, the behavior of the Earth material can be described by Hooke.s Law i.e., the

amount of strain is linearly proportional to the amount of stress. Beyond its elastic limit the material may either respond with brittle fracturing (e.g., earthquake faulting) or ductile

behavior/plastic flow.

Figure 2.1: Hook’s law showing relationship between stress and strain.

2.1 Seismic wave propagation & characteristics

The Seismic wave or Elastic wave generated at source (natural or artificial) propagates

through the medium (some portion of the earth) and recorded at receiver (called seismometer).

Seismogram (the record of the ground motion at the receiver), thus contains information about

both the source and the medium.

Precise time is incorporated in the Seismograms so that any point on the seismogram can

be related with the time. Seismograms can be related with to the actual ground motion if the

seismometer response is known. Though Seismogram appears a simply wiggly line, it contains

interesting and useful information.

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Figure 2.2: Seismogram of one day showing recording of various earthquakes.

Learning about earthquake faulting from the seismic waves that are generated by it is an

inverse problem. A major topic of seismology is the study of Seismic source –typically earthquakes. Focus or hypocenter, that is Location of an earthquake, is found from the arrival time

of seismic waves record on seismometer a different sites. The size of earthquake is measured from the amplitude of motion recorded on seismogram and given in terms of magnitude or moment.

The fault geometry, on which on an earthquake occurred is inferred from the three dimension

pattern of radiated seismic waves. The information about the amount of slip that occurred, the size

of the area that slipped and the slip process is estimated by the pulse radiated.

The seismic waves or elastic waves, which arise through sudden rupture in an earthquake

source or by an explosion, propagate through the whole of the earth’s interior or along its surface

layers. The waves are recorded by seismograph stations the world over, provided that the release

energy is big enough. When seismic energy is released suddenly at a point near the surface of a

homogenous medium, part of the energy propagates through the body of the medium as seismic

body waves. The remaining part of the seismic energy spreads out over the surface as seismic

surface wave, analogous to the ripples on the surface of a pool of water into which a stone has

been thrown.

2.1.1 Types of body waves:

There are two types of body waves:

1. P waves arrive first. Primary, compressional or longitudinal waves. Analogous to sound

waves. Particle motion is along the direction of travel (propagation) of the wave, i.e., longitudinal waves. P waves can travel through solids, liquids or gases.

2. S waves arrive second. Secondary, shear, transverse waves. S waves vibrate

perpendicular to the direction of propagation. A shear wave can be split into orthogonal,

i.e., horizontal and vertical, components. S waves do not propagate through liquids or

gases, since these don’t have any shear strength.

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Figure 2.3: Particles motion in different types of Body waves

2.1.2 Velocities of body waves:

The velocities of body waves are given dependent on the elastic constant of the media

of earth and are given by relations:

• κ= the bulk modulus (or incompressibility)

• (for granite k is about 56x1010 dynes per square cm

• for water k is about 2.0x1010 dynes per square cm )

• µ = the shear modulus (or rigidity)

• (for granite µ is about 34x1010 dynes per square cm

• for water µ =0)

• ρ = the density of the medium

• (for granite ρ is about 2.78 gm per cubic cm

• for water ρ is about 1.00 gm per cubic cm)

• For granite :Vp is about 6.0 Km/s and Vs is about 3.5 Km/s

2.1.3 Types of Surface waves:

Body-wave exist in the elastic full space. However, in the presence of a free surface, as in

the case of the Earth, other solutions are possible. They are called surface waves

There exist two types of surface waves, Rayleigh (LR or R) waves and Love (LQ or G)

waves

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2.1.3.1 Rayleigh (LR or R) waves:

Figure 2.4: Rayleigh (LR or R) waves

When a P (or SV) wave arrives at the surface the reflected wave energy contains both P

and SV waves. Lord Rayleigh showed in this case a solution of the wave equation exists for two

coupled inhomogeneous P and SV waves that propagate along the surface of a half-space. Since

Rayleigh waves originate from coupled P and SV waves they are polarized in the vertical (SV)

plane of propagation and due to the phase shift between P and SV the sense of their particle

motion at the surface is elliptical and retrograde. Rayleigh waves travel with phase velocity

slightly slower than Love waves.

2.1.3.2 Love (LQ or G) waves

Figure 2.5: Love (LQ or G) waves

Love waves only have horizontal component perpendicular to direction of propagation. Only occur when there are distinct layers. SH waves are totally reflected at the free surface.

Equivalent to trapped SH waves. Love waves are formed through constructive interference of

repeated reflections of teleseismic SH at the free-surface. The velocity of Love wave is the

velocity between the S wave velocities in the surface and deeper layers

Figure 2.6: Particles motion in different types of surface waves

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2.1.4. Effect on seismic wave paths due to internal Structure of the Earth

Due to internal structure of earth, the seismic wave path under goes the phenomenon of

reflection and refraction while traveling from the source due to different layers of the earth. P and S wave can only travel and reach the earth’s surface between 0 and 103o. Due to the

presence of liquid core only P wave is allowed to reach between 143o

and 180o. Within the

shadow zone only waves reflected from the inner core can reach the earth's surface.

Figure 2.7: Effect on seismic wave paths due to internal Structure of the Earth

2.1.5. Nomenclatures of Seismic Rays

Figure 2.8: Path of regional/teleseismic waves

P Primary wave

K P wave through outer core

I P wave through inner core

P’ Abbreviation for PKP

PP Reflected P wave with 2 legs

pP P wave with leg from focus to surface

SP S wave reflected as P wave

S Secondary wave

J S wave through inner core

SSS Reflected S wave with 3 legs

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sS S wave with leg from focus to surface

PS P wave reflected as S wave

c Wave reflected at outside boundary of outer core (e.g., ScS)

i Wave reflected at outside boundary of inner core (e.g., PKiKP)

m No. of reflections inside the outer boundary of outer core is m-1

d Depth in km from which a seismic ray is reflected

h Wave that may be reflected from a discontinuity around inner core

dif P,S Diffracted P or S waves around outer core

LQ Love waves

LR Rayleigh waves

2.2 Travel-time tables and Velocity models

Many seismologists have compiled arrival time data set shown in figure 2.9(a). The average fits

to the various families of arrival are known as the travel time curves or chart. A travel time curve

is a graph of arrival times, of various seismic waves recorded at different points as a function of

distance from the seismic source. Seismic velocities within the earth can be computed from the

slopes of the resulting curves. The tabular form of these curve are called Travel time tables.

A velocity structure is a generalized regional model of the earth's crust that represents crustal

structure using layers having different assumed seismic velocities.

Figure 2.9(a): Travel-Time picked from phase of select earthquakes and explosions with known

or well determined locations.

The first widely adopted empirical travel-time curve was published by Sir, Harold Jeffrey and Keith Bullan in 1940. The tabular form of these curve are called J-B tables.

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A plot of travel times of teleseismic rays against epicentral distance ∆ provides the basic observational data base. One of the most common travel time table which is used by

seismologist is Jeffreys-Bullen travel-time diagram for earthquake phases (1940). Surface wave plots of T vs ∆ are straight lines due to constant velocity along path. Body wave plots of

T vs ∆ are curves because velocity changes with depth.

The J-B tables are remarkable accurate, and for teleseimic distances they can predict

the travel times of principal seismic phase to with a few seconds.

Figure 2.9(b): Jeffreys-Bullen Travel-Time Diagrams;

The J-B tables, published by Jeffrey and Bullen in 1940, had been in use as standard earth

model till late 1980’s, which was later revised with more data and computer power and called as ‘iasp91’ model. The iasp91 model was developed by Kennett and Engdahl (1991) as part of

an effort of the sub-commission on ‘Earthquake Algorithms’ of the International Association of Seismology and the Physics of the Earth’s Interior (IASPEI) to generate new global travel

time tables for seismic phases. The most significant differences between the ‘iasp91’ and the older J-B travel-time model are for the upper mantle and core phases.

.

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Figure 2.10: IASP91 travel-time Diagrams (at 0 Km depth;)

The currently most common global 1-D Earth model IASP91 (Kennett and Engdahl, 1991;)

assumes a homogeneous 35 km thick two-layer crust with the intermediate crustal discontinuity at

20 km depth. The respective average velocities for the upper and lower crust and the upper mantle

are for P waves 5.8 km/s, 6.5 km/s and 8.04 km/s, and for S waves 3.36 km/s, 3.75 km/s and 4.47

km/s, respectively. The impedance contrast at the Conrad discontinuity and the Moho is about 1.3. Fig. 2.11 is a simplified depiction of such a two-layer crust and of the seismic rays of the main

crustal/upper mantle phases to be expected. These are: Pg, Sg, Pb, Sb, Pn, Sn, PmP and SmS.

Fig. 2. 11 A simplified model of the crust showing the ray traces of the main .crustal phases. observed for near (local and regional) earthquakes. Note: P* = Pb and S* = Sb.

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Subsequently, Kennett et.al., (1995) produced another model called, ‘ak135’, which gives a significantly better fit to a broad range of phases than the ‘iasp91’. The differences in velocities

between ‘ak135’ and ‘iasp91’ models are generally quite small except at the boundary of the inner core, where reduced velocity gradients are needed to achieve satisfactory performance for PKP

differential time data. The ‘ak135’ is now considered the best model for global earthquake

locations.

Figure 2.12: AK135 (seismic wave speeds according to Kennett et al. (1995), attenuation

parameters and density according to Montagner and Kennett (1996); The

abbreviation on the outermost right stand, within the marked depth ranges, for: C .

crust, UM . upper mantle, TZ . transition zone, LM . lower mantle, D''-layer, OC .

outer core, IC . inner core.

Presently, IMD employs ‘local velocity structure models’, where available, for locating local events and ‘iasp91’ model for locating regional (1,000–2,000 km) and teleseismic events (>2000

km), using Hypocenter program, which is integrated with SEISAN software for locating local, regional and teleseismic events. While the ‘Response Hydra’ autolocation software of RTSMN

system uses ‘ak135’ velocity model, the SeisComp3 software of northeast telemetry system employs ‘iasp91’ velocity model.

***********

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3 Earthquake source parameters;

Although during an earthquake, the rupture may involve area of fault-plane measuring many square kilometer in area, yet it may began from the point. The point at which the rupture

began and the first seismic waves originated is called the focus or hypocenter of the earthquake.

It generally occurs at focal depth many kilometers below the Earth surface. The point on the

Earth’s surface vertically above the focus is called the epicenter of the earthquakes.

Figure 3.1: Earthquake epicentral parameters

3.1 Magnitude, intensity, energy; etc.;

Seismologists measure the size or strength of earthquakes by their magnitude and

intensity. Prior to the development modern instrumentation, it was characterized by qualitative

description of the effect of the earthquakes. Developments of modern instrumentation have

allowed the development of various quantitative measures of earthquake size.

Magnitude is a quantitative measure of the size of the earthquake and estimates the

amount of energy released at the hypocenter. Magnitude of an earthquake is determined using the

seismograms recorded by seismographs. Intensity measures the strength of shaking produced by

the earthquake at any particular location on earth’s surface and is dependent on the distance from the epicenter. The intensity scales are primarily based on three factors – perception by people,

performance of buildings, and changes to natural surroundings. There are many intensity scales. Two commonly used ones are the Modified Mercalli Intensity (MMI) Scale and the Medvedev-

Sponheuer-Karnik (MSK) Scale. Both scales are quite similar and range from I (least perceptive) to XII (most severe).

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3.1.1 Earthquake magnitudes

The more objective and quantitative measure of the size of an earthquake came after the development of modern instrumentation, which is based on the amplitude of the resulting wave

amplitude recorded on the seismogram. The concept of the wave amplitude reflect the earthquake

size once the amplitude are corrected for the decreasing with the distance due to geometric

spreading and attenuations.

Magnitude scale have been have the general form;

M=log (A/T) + f (h, ∆ ) + C

Where A is the amplitude of the signal,

T is the dominant period,

f is a correction for the variation of amplitude with the earthquake’s depth h and distance

∆ from the seismometer.

and C is a regional scale factor.

3.1.1.1 Local Magnitude (ML):

The earliest magnitude scale , introduced by Charles Richter in 1935 is the local

magnitude, ML, often referred as the “Richter Scale’. ML is determined from the maximum amplitude measured in microns on a seismogram written by Wood-Anderson seismograph with

free period of 0.8 second, magnification of 2,800, damping factor of 0.8 calculated to be at a distance of 100 kms. The amplitude of the largest arrival (often the S wave) is measured and

corrected for the distance between the source and the receiver, given by the difference in the arrival times of P and S waves.

The scale ML =log A – 2.48 + 2.76 * log ∆ is for instrument period (0.8S) and the

distances in Km

ML was originally proposed to be used for distances upto 600 kms and is a very useful

scale for engineering applications. Many structures have natural periods close to that of Wood -

Anderson instrument (0.8 seconds) and the extent of earthquake damage is closely related to ML.

Figure 3.2: Nomogram for Richter magnitude.

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3.1.1.2 Surface wave magnitude (Ms):

For larger epicentral distances, say, beyond about 1000 Km. the seismogram is usually

dominated by surface waves with longer periods. Surface waves contain the highest energy

content of the whole wave front. It is thus possible to estimate the energy released by an

earthquake from the amplitudes and periods of surface waves.

The surface wave magnitude is computed using the formula,

Ms = log (A/T) +1.66 log ∆ + 3.3

Where, A is the maximum ground amplitude (in micron) of surface (Rayleigh) waves of

period 20+2 sec, T is the corresponding period of the wave and

∆ the epicentral distance in geocentric degrees.

The surface wave magnitude is most commonly used to describe the size of shallow,

distant moderate to large earthquakes

3.1.1.3 Body wave magnitude (mb):

mb is measured for the early portion of the body wave train, usually for the P wave , using

mb = log (A/T) + Q (h, ∆ )

where A is the ground motion amplitude (in microns) after the effects of the seismometer

are removed.

T is corresponding wave period (in seconds) and

Q (h, ∆) is a empirical correction factor which depends on focal depth h and epicenter

distance ∆.

mb = log A -log T + 0.01 ∆ -5.9

Common practice is to use the first 5-10 cycles of the records and period of about 1s on

instrument with peak response near 1S . mb is measured out to 100 0 distances, beyond which

diffractions around the core has a complicated effect on amplitudes.

3.1.1.4 Moment magnitude (Mw):

Magnitude scale based on seismic moment (M0) was proposed by Kanamori (1977). The Moment magnitude (Mw) is estimated using the formula given below:

Mw = (log M0-16)/1.5

Where M0 is the seismic moment in dyne-cm.

The static seismic moment is a product of rupture area, shear strength and fault

displacement. M0 = µD A

with µ - rigidity or shear modulus of the medium, D - average final displacement after the

rupture, A - the surface area of the rupture.

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Seismic moment can also be estimated from the displacement spectra of ground motion time histories or waveform modeling/ inversion. Since seismic moment is a measure of strain

energy released from the entire rupture surface, a magnitude scale based on seismic moment most accurately describes the size of the largest earthquakes. Since M0 does not saturate, so also Mw.

3.1.2 Earthquake intensity:

Large earthquakes produce alterations to the Earth’s natural surface features, or severe

damage to the man-made structures such as buildings, bridges and dams. Even small earthquakes can result in disproportionate damage to these edifices when inferior constructions methods or

materials have been utilized. The intensity of earthquake at a particular place is classified on the basis of the local character of the visible effect it produces. Have advantages that it is inferred

from human account, so can be determined where no seismometer was present. Intensity is often give the best information available about the historical earthquakes.

Various types of scales have been developed for the classification of intensity. The Rossi-

Forel (RF) scale of intensity (I – X), was developed in 1880 and used for many years. It has been

replaced by Modified Mercalli Intensity (MMI) scale, which was originally developed by seismologist, G Mercalli. Uses roman numeral ranging from I (generally unfelt) To XII (total

destructive). The Japanese Meteorological Agency (JMA) has its own intensity scale, and the Medvedev-Spoonheuer-Karnik (MSK) scale is also now widely used.

Table 1: MODIFIED MERCALLI INTENSITY SCALE (ABRIDGED)

Intensity EEffffeeccttss

I Not felt except by a very few under especially favorable circumstances.

II Felt only by a few persons at rest, especially on upper floors of buildings. Delicately

suspended objects may swing

III Felt quite noticeably indoors, especially on upper floors of buildings, but many people do

not recognize it as an earthquake. Standing automobiles may rock slightly. Vibrations like

passing of truck. Duration estimated.

IV During the day felt indoors by many, outdoors by few. At night some awakened. Dishes,

windows, doors disturbed; walls make creaking sound. Sensation like heavy truck striking

building. Standing automobiles rocked noticeably. V Felt by nearly everyone, many awakened. Some dishes, windows and so on broken; cracked

plaster in a few places; unstable objects overturned. Disturbances of trees and poles, and

other tall objects sometimes noticed. Pendulum clocks may stop.

VI Felt by all, many frightened and run outdoors. Some heavy furniture moved; a few instances

of fallen plaster and damage chimneys. Damage slight.

VII Everybody runs outdoors. Damage negligible in buildings of good design and construction;

slight to moderate in well–built ordinary structures; considerable in poorly build or badly

designed structures; some chimneys broken. Noticed by persons driving cars.

VIII Damage slight in specially designed structures; considerable in ordinary substantial

buildings with partial collapse; great in poorly built structure. Panel walls thrown out of

frame structures. Fall of chimneys, factory stacks, columns, monuments, and walls. Heavy

furniture overturned. Sand and mud ejected in small amounts. Change in well water. Persons driving cars disturbed.

IX Damage considerable in specially designed structures; well-designed frame structures

thrown out of plumb; great in substantial buildings, with partial collapse. Buildings shifted

off foundations. Ground cracked conspicuously. Under Ground pipes broken.

X Some will-built wooden structures destroyed; most masonry and frame structures destroyed

with foundations; ground badly cracked. Rail bent. Land slides considerable from

riverbanks and steep slopes. Shifted sand and mud. Water splashed, slopped over banks.

XI Few, if any, (masonry) structures remain standing. Bridges destroyed. Broad fissures in

ground. Underground pipelines completely out of service. Earth slumps and land slips in

soft ground. Rail bent greatly.

XII Damage total. Waves seen on ground surfaces. Lines of sight and level distorted. Objects thrown into the air.

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3.1.2.1 Isoseismals:

The variation of intensities with distance from an earthquake can be seen by plotting line

of constant intensity know as Isoseismals.

Figure 3.3: Isoseismal map for Bhuj earthquake of 26th January, 2001

Variation in the acceleration with magnitude and distance for the earthquake can be

described approximately by relation like:

a (M,r)=b 10 c M

r –d

Where b,c and d are constants that depend on factors including: geology of the area, the

earthquake depth and fault geometry and the frequency of ground motion.

3.1.3 Relation between magnitude and energy:

The definition of earthquake magnitude relates it to the logarithm of the amplitude of a

seismic disturbance. As the energy of a wave is also proportional to the square of its amplitudes

therefore the magnitude is also related to the logarithm of the energy.

An empirical relation worked out by Gutenberg and Richter, relates the energy release E to the surface-wave magnitude Ms

Log 10 E= 4.4 +1.5 Ms where E is in Joules.

Log 10 E= 11.4 +1.5 Ms where E is in Ergs.

An alternate version of the energy-magnitude relation, suggested by Bath for magnitudes

Ms>5.0 is Log 10 E= 5.24 +1.44 Ms

The logarithmic nature of each formula means that the energy release increases very

rapidly with increasing magnitude. A unit change in magnitude corresponds to 32 fold increase in

seismic energy.

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3.2. Earthquake statistics;

3.2.1 Magnitude –frequency relationship:

Every year there are many small earthquakes, and only a few large ones. Gutenberg and Richter noted that this trend of magnitude –frequency relationship appears to obey a power law

and can be given by an empirical formula: log N = a-b M , Where N is the number of events with

magnitudes in the range M + ∆M.

Figure 3.4: Worldwide earthquake frequency-magnitude plot

The values of a varies between about 8 and 9 from one region to another, while b known as b-value is generally lies between 0.8 and 1.2 for a wide variety of region.

3.3. Digital data analysis and location of earthquakes;

Continuous digital waveform data generated by digital seismograph systems is extensively

used in modern seismological research, particularly for studying the source, site and path

characteristics of the media.

3.3.1 Determination of epicenter of an earthquake;

One of the most important tasks in observation seismology is locating seismic sources. This involves determining both the hypocentral coordinated and the source origin time; in general

determination of the source location requires identification of seismic phases and measuring their arrival time, as well as knowing the velocity structure between the hypocenter and seismic station.

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Figure 3.5: Steps in generation of seismic information

3.3.1.1 Multiple station location.

To locate an earthquake, one need the data from at least three seismometer stations. The

seismometer records the time when the P and S-waves arrive at the recording station. P-waves

travel faster through the earth than S-waves and so they arrive at the seismometer station before

the S-waves and are recorded by the seismometer first. The difference in arrival time between the

two types of seismic wave can be used to calculate the distance of the earthquake's epicentre from the seismometer, as the further away an earthquake is, the greater the lag time between the

detection of the S waves relative to the P waves. Based on properties of the crust, and many trials, a seismologist can calculate how far away an earthquake is from a station based just on the S-P

lag time.

Figure 3.6: Location by the circle (or arc) method

Since P wave travels faster than S wave, P wave will arrive first on the seismogram. The

difference between the arrival time will depend on the p wave and s wave velocities.

D= ∆tp-s / (1/VS -1/Vp)

The time separation, ∆tp-s between the P- and S-wave arrivals multiplied by the ratio

VP·VS/ (VP-VS) of the P- and S-wave velocities gives us the epicentral distance (distance from the

station to the projection of the earthquake's focus at the surface). The epicenter is found within the

black area where the circles cross.

These circles will rarely cross at one point, which indicates errors in the observations,

errors in the model, and/or a subsurface depth. With only two stations, we see that there are either

two possible locations, or no possible location if the two circles do not intersect. With more than

three stations, the uncertainty in location decreases

Earthquake Seismograph Estimation of epicenter

Products

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The accuracy of these techniques will depend on the number, quality, geographic

distribution of these seismographs and the on the accuracy of seismic velocity model.

Earlier, these above exercises were done manually which used to takes a lot of time. Now

with availability of sophisticated computational techniques, various software are available to analyze the data and determined the locations and other parameter very fast and thus reduced the

time response in the determination of epicentral and other source parameters an earthquake.

3.3.1.2 Automatic Earthquake Location

Finding the earthquake location automatic is generally posed as an inverse problem. In this

we know the phase or arrival time of the event and then we solve it for a source location and

origin time which are consistent with our recorded data.

Figure 3.7: Auto Location of earthquake with inverse method.

• P-wave only method.

• The location has 4 unknowns (t,x,y,z) so with 4+ P arrivals this can be solved.

• The P arrival time has a non-linear relationship to the location,– therefore can only

be solved numerically

• Use a Least squares method – minimize residuals between observed and calculated

travel times

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3.4. Seismological operations and information dissemination.

The science of Seismology has not yet reached a stage, anywhere in the world, to be able to

predict the occurrence of earthquakes with reasonable degree of accuracy with regard to space, time and magnitude. It is, therefore, not possible to issue any warnings, in advance, regarding

the occurrence of future earthquakes. However, timely dissemination of earthquake

information, to the concerned state and central government agencies, plays a very important

role in dealing with the post-disaster relief and rehabilitation related matters leading to mitigation

of the losses. Also, loss of life and property caused due to earthquakes may be reduced

substantially, through proper planning and implementation of pre- and post-disaster management

strategies. Towards achieving these objectives, India Meteorological Department (IMD) is

engaged in generating and providing the required Science and Technology (S&T) related inputs,

to all the concerned state and central government organizations and other user agencies.

IMD maintains round-the-clock monitoring of seismic activity in and around the country.

The operational task of the department is to quickly estimate the source parameters, immediately

after the occurrence, of earthquakes occurring in and around the country and disseminate the

information to all concerned agencies. Towards meeting the above objectives, India Meteorological Department is maintaining the national seismological network consisting of a total

of 82 seismological stations, spread over the entire length and breadth of the country. This includes: a) 16-station V-SAT based digital seismic telemetry system around National Capital

Territory (NCT) of Delhi, b) 20-station VSAT based real time seismic monitoring network in North East region of the country and (c) 17-station Real Time Seismic Monitoring Network

(RTSMN) to monitor and report large magnitude under-sea earthquakes capable of generating tsunamis on the Indian coastal regions and rest are standalone/analog.

A Control Room is in operation, on 24X7 basis, at IMD Headquarters (Seismology) in New

Delhi with state-of-art facilities for data collection, processing and dissemination of information to

the concerned user agencies. As part of early warning system for tsunamis established by the Indian National Center for Ocean Information Services (INCOIS), Hyderabad, a 17-station Real

Time Seismic Monitoring Network (RTSMN) was set up by IMD to monitor and report large magnitude under-sea earthquakes capable of generating tsunamis on the Indian coastal regions.

The ground motion data recorded at the 17 field stations is transmitted in real time through VSAT communication systems to the two Central Receiving Stations (CRSs) located at IMD, New Delhi

and INCOIS, Hyderabad for processing. The RTSMN system employs state-of-art auto-location software, called ‘Response Hydra’ (v-1.2), to make preliminary estimates of earthquake source

parameters immediately (within a few minutes) after the occurrence of an earthquake. The source

parameters include the time of occurrence, location (region), magnitude and focal depth of the

earthquake. These estimates are refined and finalized as more and more data becomes available

subsequently. The earthquake information is disseminated to the concerned user agencies through

various modes of communications, such as SMS, FAX, Email and IVRS and also posted on

IMD’s Website.

The real time seismic monitoring network established by IMD in northeast India employs

SeisComp3 software for auto-location of the earthquake events. SeisComp3 is one of the most

widely distributed free-ware software packages, designed as a high standard fully automatic data

acquisition and (near-) real-time data processing tool including quality control, event detection

and location as well as dissemination of event alerts.

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Figure 3.8: National Seismological Network

Figure 3.9: Delhi Telemetry Network Figure 3.10: Northeast Telemetry Network

As per the detection and location capabilities of the existing seismological network (s)

being operated by IMD, earthquakes of magnitude: M:~3.5 & above in Peninsular region, M:~4.0 & above in the extra-Peninsular region, M:5.0 & above in the border regions, M:6.5 & above

capable of generating Tsunamis on Indian Coasts / Territories and M:~2.0-2.5 in and around

Delhi, are reported in operational mode. In case of low intensity tremors reported felt anywhere

within the country and / or not auto-located by the RTSMN system, the data is processed

manually and earthquake information is transmitted to the concerned local authorities.

Immediately on occurrence of a locatable earthquake, the RTSMN system makes automatically a

preliminary estimate of the earthquake source parameters within a few minutes and displays the

NDI

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same on the plasma monitor in CRS. Following is the sequence of actions that follow subsequently:

1. The preliminary earthquake information is transmitted automatically through an SMS

message (Level-I) and an email message to a few designated authorities / officials of IMD

& MoES. The earthquake information includes the time of occurrence of the earthquake,

it’s location (epicenter and region), magnitude and focal depth. These estimates are refined

automatically by the system as and when more data becomes available subsequently in real

time.

2. The Duty Officer at the Control Room scrutinizes the preliminary estimates and seeks

approval of the competent authority for dissemination to various user agencies through

different modes of communication as detailed below:

i. Issue of SMS (Level-II) message of scrutinized estimates to the designated

authorities / officials..

ii. Dissemination of Preliminary Earthquake Report (PER) through FAX

to the designated authorities / officials. iii. Check and ensure updation of IMD’s Website (www.imd.gov.in), for

inclusion of the information on the current event. iv. Ensure dissemination of earthquake email message based on the revised

estimates, to the designated authorities / officials. Attach MT / CMT

solutions (if available, for large events).

v. Ensure compliance of guidelines to be followed for the issue of Preliminary Earthquake Report (PER) to the Integrated Operation Centre (IOC),

MHA. vi. Keep a log / record of important telephone calls / messages received,

particularly in connection with the occurrence of significant earthquakes.

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