Presentation Transcript

1. PSTE 4223 Methodessismiques

   Part I: Seismic Refraction

   Anne Obermann 2 x 3h
2. Overview Introduction – historical outline Chapter 1: Fundamental concepts Chapter 2: Data acquisition and material Chapter 3: Data interpretation A: Geophysical Interpretation B : Geological Interpretation
3. Overview Introduction – historical outline Chapter 1: Fundamental concepts Chapter 2: Data acquisition and material Chapter 3: Data interpretation
4. Towards refraction seismology 1885 all that was known about the Earth structure was a vague idea that the density inside had to be much greater than at the surface within 50 years an incredible amount more had been learned using seismology Breakthrough: Seismometer (late 1800’) Instrumental challenge: how to measure ground motion given that the seismometer sits on the ground? Record very small ground motions on the order of 10-3 cm for distantearthquakes
5. Towards refraction seismology Seismometers were developed to record vertical and horizontal motion. Precise timing, nowadays done using GPS (Global Positioning System) clocks - so that records can be compared between stations. Data are now recorded digitally and made available on the web.
6. Towards refraction seismology In 1889, an earthquake in Japan was recorded successfully on several seismometers in Germany. Milne discovered that observations showed that the time separations between P and S wave arrivals increased with distance from the earthquake. Thus, the S-P time could be used to measure the distance to the earthquake.
7. Towards refraction seismology Next step: Infer the velocity structure of the Earth as a function of depth from the seismograms that were recorded from many different earthquakes (Inverse Problem). The simplest approach to the inverse problem treats the earth as flat layers of uniform velocity material. The basic geometry is a layer of thickness z, with velocity v1, overlying a halfspace with a higher velocity v2. Problem: Data quantity dependent on large earthquakes – different sources needed!
8. Towards refraction seismology
9. Towards refraction seismology Set out a line or array of geophones Input a pulse of energy into the ground Record the arrival times to interpret the velocity structure
10. Towards refraction seismology
11. Seismic methods and scales Controlled source seismology - allows higher resolution studies (m to 100s km) - can carry out experiments away from tectonic regions Global seismology (earthquakes) - provides information on global earth structure and large scale velocity anomalies (100s to 1000s km) - difficult to image smaller scale structure, particularly away from earthquake source regions
12. Seismic methods and scales Seismic refraction - Used to study large scale crustal layering: thickness and velocity Seismic reflection - “Imaging” of subsurface reflectors - Difficult to determine accurate velocities and depths Refraction Reflection
13. Applications
14. Overview Introduction – historical outline Chapter 1: Fundamental concepts - Physical notions - Two-layered model - Special cases Chapter 2: Data acquisition and material Chapter 3: Data interpretation
15. Different waves P (compression) + S (shear) waves Surface waves
16. Huygens Principle Each point along a material acts like a point source of waves. Waves have circular (spherical) wave fronts, these interact constructively (destructively) and produce the wave fronts that we plot as rays.
17. Snell’s Law Wrong representation !! Why? -> Seismic rays obey Snell’s law The angle of incidence equals the angle of reflection. The angle of transmission is related to the angle of incidence through the velocity ratio. Note: the transmitted energy is refracted
18. Snell’s law: S wave conversion A conversion from P to S or vice versa can also occur. Still, the angles are determined by the velocity ratios. α1, β1 p is the ray parameter and is constant along each ray. α2, β2
19. Snell’s law: Critical Incidence whenα2 > α1, e2 > i =>we can increase iP until e2 = 90° when e2=90 °, i=ic the critical angle α1 The critically refracted energy travels along the velocity interface at α2 continually refracting energy back into the upper medium at an angle ic. α2 Head wave
20. Wave Propagation according to Huygens Principle Wrong representation !!! ->
21. Wave Propagation according to Huygens Principle
22. Wave Propagation according to Huygens Principle
23. Wave Propagation according to Huygens Principle
24. Seismic Method comparison
25. Seismic Method comparison
26. Two-layered model
27. Two-layered model Energy from the source can reach the receiver via different paths Direct wave Reflected wave Head wave or Refracted wave
28. Time-Distance Diagram (Travel Time curves) Think about: What would a fast velocity look like on this plot? Why is the direct ray a straight line? Why must the direct ray plot start at the origin (0,0)? Why is the refracted ray a straight line? Why does the refracted ray not start at the origin? Why does the reflected ray start at origin? Why is the reflected ray asymptotic with the direct ray?
29. Two-layered model Slope=1/v1 1. Direct wave Energy travelling through the top layer, travel-time Time (t) The travel-time curve for the direct wave is simply a linear function of the seismic velocity and the shot-point to receiver distance Distance (x) Receiver Shot Point Direct Ray v1 x
30. Two-layered model 1. Direct wave 2. Reflected wave Energy reflecting off the velocity interface. As the angles of incidence and reflection are equal, the wave reflects halfway between source and receiver. The reflected ray arrival time is never a first arrival. Time (t) Distance (x) Receiver Shot Point v1 Layer 1 v2 Layer 2
31. 2. Reflected wave The travel time curve can be found by noting that x/2 and h0 form two sides of a right triangle, so This curve is a hyperbola, it can be written as For x = 0 the reflected wave goes straight up and down, with a travel time of TR(0) = 2h1/v1. At distances much greater than the layer thickness (x >> h), the travel time for the reflected wave asymptotically approaches that of the direct wave. Time (t) “Intercept Time” Gives Layer Thickness Distance (x) Receiver Shot Point x h1 v1 Layer 1 v2 Layer 2
32. Two-layered model 1. Direct wave 2. Reflected wave 3. Head wave or Refracted wave -Energy refracting across the interface. -Only arrives after critical distance. - Is first arrival only after cross over distance critical distance cross over distance Time (t) Distance (x) “Critical Distance” No Refracted Rays  ic ic ic ic v1 Layer 1 Layer 2 v2
33. Reminder 3. Head wave or Refracted wave The travel time can be computed by assuming that the wave travels down to the interface such that it impinges at critical angle, then travels just below the interface with the velocity of the lower medium, and finally leaves the interface at the critical angle and travels upwards to the surface. Show that: . X x0 v1 D ic ic A h1 v2 B C
34. 3. Head wave or Refracted wave The axis intercept timeis found by projecting the travel time curve back to x = 0. The intercept time allows a depth estimation. Critical distance xc: distance beyond which critical incidence first occurs. Atthecriticaldistancethe direct wave arrives before the head wave. At some point, however, the travel time curves cross, and beyond this point the head wave is the first arrival. The crossover distance, xd, where this occurs, is found by setting TD(x) = TH(x) , which yields: The crossover distance is of interest to determine the length of the refraction line.
35. Travel-time for refracted waves critical distance cross over distance Time (t) Distance (x)
36. Note on Refraction angle Reminder: Interestingtonoticethatthehigherthevelocitycontrast, thesmallertherefraction angle. V1 = 1000 m/s λ = 11 ° V2 = 5000 m/s V1 = 1000 m/s λ = 30 ° V2 = 2000 m/s => We can only analyse cases with an increasing velocity function with depth
37. Summary h1 v1 determined from the slope of the direct arrival (straight line passing throughtheorigin) v2 determined from the slope of the head wave (straight line first arrival beyondthecriticaldistance) Layer thickness h1 determined from the intercept time of the head wave (already knowing v1 and v2)
38. Multiple-layers For multiple layered models we can apply the same process to determinelayerthicknessandvelocity sequentially from the top layertothebottom.
39. Multiple-layers The layer thicknesses are not as easy to find Recall… Solve for h1… Now, plug in h1 and solve the remaining layers one at a time … BEWARE!!! h1, h2, are layer thicknesses, not depth to interfaces. So, depth to bottom of layer 3 /top of layer 4 = h1 + h2 + h3
40. Multiple-layers General formulation
41. Overview Introduction Chapter 1: Fundamental concepts Chapter 2: Material and data acquisition Chapter 3: Data interpretation
42. Material Geophones Recordingdevice (Computer, Seismograph) Source (hammer, explosives) Battery Cables (Geode)
43. Material: Geophones Geophones need a good connection to the ground to decrease the S/N ratio (can be buried)
44. Material: Cable, Geode
45. Material: Energy Source Sledgehammer (Easy to use, cheap) Buffalo gun (More energy) Explosives (Much more energy, licence required) Drop weight (Need a flat area) Vibrator (Uncommun use for refraction) Air gun (For lake / marine prospection) Goal: Produce a good energywithhighfrequencies, Possible investigation depth 10-50 m You canadd (stack) few shots to improve signal/noise ratio
46. Data acquisition Number of receivers and spacing between them => will define length of the profile and resolution Number of shots to stack (signal to noise ratio) Position of shots
47. Geophone Spacing / Resolution Often near surface layers have very low velocities E.g. soil, subsoil, weathered top layers of rock These layers are likely of little interest, but due to low velocities, time spent in them may be significant To correctly interpret data these layers must be detected Find compromise between: Geophone array length needs to be 4-5 times longer than investigation depth Geophone distance cannot be too large, as thin layer won’t be detected
48. Geophone Spacing / Resolution This problem is an example of…?
49. Overview Introduction Chapter 1: Fundamental concepts Chapter 2: Data acquisition and material Chapter 3: Data processing and interpretation
50. Record example Dynamite shot recorded using a 120-channel recording spread
51. Record example Example of seismic refraction data acquisition where students are using a 'weight-drop' - a 37 kg ball dropped on hard ground from a height of 3 meter - to image the ground to a depth of 1 km
52. Record example Time Time Distances
53. First Break Picking This is the most important operation, good picking on good data !!!! A commun problemis the lack of energy, for far offset geophones
54. First Break Picking –on good data noise
55. First Break Picking –on poor data noise ?
56. Travel-time curve How does the inverse shot look like in an planar layered medium? t distance
57. Reciprocity of travel-times
58. Assigning different layers
59. Control of travel-times
60. Travel time inversion to find best matching underground model
61. Complete analysis process
62. Exercice
63. Some Problems Dipping interfaces Undulating interfaces There are two cases where a seismic interface will not be revealed by a refraction survey. The low velocity layer The hidden layer
64. Dipping Interfaces What if the critically refracted interface is not horizontal? A dipping interface produces a pattern that looks just like a horizontal interface! Velocities are called “apparent velocities” What do we do? In this case, velocity of lower layer is underestimated underestimated
65. Dipping Interfaces To determine if interfaces are dipping… Shoot lines forward and reversed If dip is small (< 5o) you can take average slope The intercepts will be different at both ends Implies different thickness Beware: the calculated thicknesses will be perpendicular to the interface, not vertical
66. Dipping Interfaces If you shoot down-dip Slopes on t-x diagram are too steep Underestimates velocity May underestimate layer thickness Converse is true if you shoot up-dip In both cases the calculated direct ray velocity is the same. The intercepts tint will also be different at both ends of survey
67. Problem 1: Low velocity layer If a layer has a lower velocity than the one above… There can be no critical refraction - The refracted rays are bent towards the normal There will be no refracted segment on the t-x diagram for the second layer The t-x diagram to the right will be interpreted as - Two layers - Depth to layer 3 and thickness of layer1 will be exaggerated Causes: Sand below clay Sedimentary rock below igneous rock (sometimes) sandstone below limestone How Can you Know?
68. Problem 2: Hidden layer Recall that the refracted ray eventually overtakes the direct ray (cross over distance). The second refracted ray may overtake the direct ray first if: The second layer is thin The third layer has a much faster velocity
69. Undulating Interfaces Undulating interfaces produce non-linear t-x diagrams There are techniques that can deal with this delay times & plus minus method We will see them later…
70. Detecting Offsets Offsets are detected as discontinuities in the t-x diagram Offset because the interface is deeper and D’E’ receives no refracted rays.
71. Question: To which type of underground model correspond the following travel-time curves? t t distance distance
72. Further information http://www.geomatrix.co.uk/training-videos-seismic.php