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Reflection seismology 1

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  • 1. Seismic wave propagation: Connected set of disturbance and wave motion is perpendicular to these wave fronts 3
  • 2. When do we have a good reflection ??? 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. 4
  • 3. In practice, reflection seismology is carried out at comparatively small angles of incidence 5 More reflection coefficient better reflection.
  • 4. 6
  • 5. • 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 7
  • 6. Reflected P- and S-waves and refracted P- and S-waves are generated from the incident P-wave.
  • 7. CRITICAL ANGLE AND HEAD WAVES Critical Distance Critical angle Vp Lower Vp Higher The passage of the refracted wave along the Angle of refraction= 90 interface in the lower medium generates a plane wave traveling upward in the upper medium. Critically refracted waves are called head waves
  • 8. 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.
  • 9. 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.
  • 10. 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
  • 11. x S G θ 2d R Vt V= velocity T= travel time S’ Hence in right angled triangle SS’G we have x2 + (2d)2 = (vt)2 Equation of Hyperbola {(vt)2 / 4d2 }-{(x2)/4d2}=1 Hence, reflection travel time curve are hyperbola.
  • 12. 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 traveltime 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.
  • 13. • This can be determined from t0, the two-way reflection traveltime 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 to approximately. And we have: Or, As d= Vt0 equ. 1 (Higher order terms have been neglected here
  • 14. 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.