3

Connected set of

disturbance and wave

motion is perpendicular to

these wave fronts

Seismic wave propagation:

4

Acoustic impedance: The acoustic impedance of a

rock is determined by multiplying its density by its

seismic wave-wave velocity, i.e., V. Acoustic impedance

is generally designated as Z.

When do we have a good reflection ???

5

More reflection coefficient better reflection.

In practice, reflection seismology is

carried out at comparatively small

angles of incidence

6

7

• For SMALL INCIDENT ANGLES, all of the energy is

the reflected or transmitted Pwaves there are

essentially no S-waves.

• As the incident angle increases some of the energy

goes into reflected and transmitted S-waves

Conversion of P

wave in to s wave

at the interface

Reflected P- and S-waves and refracted P- and S-waves are

generated from the incident P-wave.

CRITICAL ANGLE AND HEAD WAVES

Vp

Higher

Critical

angle

Angle of refraction= 90

Critical

Distance

Critically refracted waves are called head waves

Vp

Lower

The passage of the refracted wave along the

interface in the lower medium generates a plane

wave traveling upward in the upper medium.

Subcritical reflection: Angle of incidence less than the critical angle.

Critical reflection: The ray that is incident on the boundary at C is

called the critical ray because it experiences critical refraction. The critical

ray is accompanied by a critical reflection. It reaches the surface at a critical

distance (xc) from the source at O.

Supercritical reflection: The seismic rays that are incident more

obliquely than the critical angle are reflected almost completely. These

reflections are termed supercritical reflections, or simply wide-angle

reflections.

REFLECTION SEISMOLOGY

� finding the depths to reflecting surfaces and the seismic velocities of subsurface rock layers

� Principles:

1. A seismic signal (e.g., an explosion) is produced at a

known place at a known time, and the echoes reflected

from the boundaries between rock layers with

different seismic velocities and densities are recorded

and analyzed.

2. Compactly designed, robust, electromagnetic

seismometers – called “geophones” in industrial usage

– are spread in the region of subcritical reflection,

within the critical distance from the shot-point, where

no refracted arrivals are possible.

Reflection seismic data are most usually acquired along profiles

that cross geological structures as nearly as possible normal to the

strike of the structure.

Reflection seismic data are most usually acquired along profiles

that cross geological structures as nearly as possible normal to the

strike of the structure.

Reflection at a horizontal

interface:

• d- Depth of the reflector

below the shot point.

• x- Horizontal distance from

the shot point to receiver at G

• The first signal received at G

is from the direct wave that

travels directly along SG

(body wave).

The travel-time t of the

reflected ray SRG is

(SR+RG)/V. However, SR and

RG are equal and therefore

2d

S

R

G

S’

θ

x

Vt

V= velocity

T= travel time

Hence in right angled

triangle SS’G we have

x2 + (2d)2 = (vt)2

{(vt)2 / 4d2 }-{(x2)/4d2}=1

Equation of Hyperbola

Hence, reflection travel time curve are hyperbola.

T= Reflection travel time

• At t=0, t=to (vertical travel

time given by 2d/v)

• For large distances from the

shot-point (x>>2d) the travel-

time of the reflected ray

approaches the travel-time of

the direct ray and the

hyperbola is asymptotic to the

two lines tx/V

A principle goal of seismic reflection

profiling is usually to find the vertical

distance (d) to a reflecting interface.

• This can be determined from t0, the two-way reflection travel-

time recorded by a geophone at the shot-point, once the velocity

V is known.

Determination of the velocity:

• One way of determining the velocity is by comparing t0 with

the travel-time tx to a geophone at distance x.

1). laid out geophone close to the shot-point and the

assumption is made that the geophone distance is much less

than the depth of the reflector (x<<d). This can give us toapproximately.

And we have:

Or,

equ. 1 (Higher order

terms have been

neglected here

As d= Vt0

The difference between the travel-time tx and the shotpoint

travel-time t0 is the normal moveout, ∆tn = tx – t0. By rearranging

Eq. 1 we get

The echo time t0 and the normal moveout time ∆tn are found

from the reflection data. The distance x of the geophone from the

shot-point is known and therefore the layer velocity V can be

determined. The depth d of the reflecting horizon can then be

found by using the formula for the echo time.