reflection seismology 1
TRANSCRIPT
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Connected set of
disturbance and wave
motion is perpendicular to
these wave fronts
Seismic wave propagation:
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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 ???
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More reflection coefficient better reflection.
In practice, reflection seismology is
carried out at comparatively small
angles of incidence
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• 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.