1. Stress:
Force per unit area across an arbitrary plane

2. Stress Defined as a Vector^
N = unit vector normal to plane
t(n) = (tx,ty,tz) = traction vector ^
The part of t that is perpendicular to the plane is normal stress
The part of t that is parallel to the plane is shear stress

3. Stress Defined as a Tensorz
txz y x z
tzx ^ ^
t(y)
t(z) x
t(x) ^
No net rotation
txx txy txz
t = t T = tyx tyy tyz
tzx tzy tzz

4. Relation Between the Traction Vector and the Stress Tensorz
txz y x z
tzx ^ ^
t(y)
t(z) x
t(x) ^
No net rotation ^
tx(n) txx txy txz nx
t(n) = t n = ty(n) = tyx tyy tyz ny
tz(n) tzx tzy tzz nz ^ ^ ^ ^ ^ ^ ^

5. Relation Between the Traction Vector and the Stress Tensor
That is, the stress tensor is the linear operator that produces the traction vector from the normal unit vector…. ^
tx(n) txx txy txz nx
t(n) = t n = ty(n) = tyx tyy tyz ny
tz(n) tzx tzy tzz nz ^ ^ ^ ^ ^ ^ ^

6. Principal Stresses
Most surfaces has both normal and tangential (shear) traction components.
However, some surfaces are oriented so that the shear traction = 0.
These surfaces are characterized by their normal vector, called principal stress axes
The normal stress on these surfaces are called principal stresses
Principal stresses are important for source mechanisms

7. Stresses in a Fluid -P 0 0 t = 0 -P 0 0 0 -P
If t1=t2=t3, the stress field is hydrostatic, and no shear stress exists -P
t = P P
P is the pressure

8. Pressure inside the Earth
Stress has units of force per area:
1 pascal (Pa) = 1 N/m^2
1 bar = 10^5 Pa
1 kbar = 10^8 Pa = 100 MPa
1 Mbar = 10^11 Pa = 100 GPa
Hydrostatic pressures in the Earth are on the order of GPa
Shear stresses in the crust are on the order of MPa

9. Pressure inside the Earth
At depths > a few km, lithostatic stress is assumed, meaning that the normal stresses are equal to minus the pressure (since pressure causes compression) of the overlying material and the deviatoric stresses are 0.
The weight of the overlying material can be estimated as rgz, where r is the density, g is the acceleration of gravity, and z is the height of the overlying material.
For example, the pressure at a depth of 3 km beneath of rock with average density of 3,000 kg/m^3 is
P = 3,000 x 9.8 x 3,000 ~ ^7 Pa ~ 100 MPa ~ 0.9 kbar

10. Mean (M) and Deviatoric (D) Stress
txx txy txz
t = tyx tyy tyz
tzx tzy tzz
M = txx + tyy + tzz = tii/3
txx-M txy txz
D = tyx tyy-M tyz
tzx tzy tzz-M

11. Strain:
Measure of relative changes in position (as opposed to absolute changes measured by the displacement)
U(ro)=r-ro
E.g., 1% extensional strain of a 100m long string
Creates displacements of 0-1 m along string

12. J can be divided up into strain (e) and rotation (Ω)
is the strain tensor (eij=eji)
ux ux uy ux uz
x y x z x
½( )
½( )
e =
uy ux uy uy uz x y y z y
½( )
½( )
uz ux uz uy uz
x z y z z
½( )
½( )

13. J can be divided up into strain (e) and rotation (Ω)
ux uy ux uz
y x z x
0 ½( )
½( ) Ω=
uy ux uy uz x y z y
-½( )
½( )
uz ux uz uy
x z y z
-½( )
-½( )
is the rotation tensor (Ωij=-Ωji)

14. Volume change (dilatation)
= 1/3 ( ) = tr(e) = div(u)
> 0 means volume increase
< 0 means volume decrease
ux uy uz
x y z
ux x
ux x
> <0

15. ∂2ui/∂t2 = ∂jij + fi = Equation of motion = Homogeneous eom when fi=0

16. Seismic Wave Equation (one version)
For (discrete) homogeneous media and ray theoretical methods, we have 
∂2ui/∂t2 = ()·u-xx u

17. ^ Plane Waves: Wave propagates in a single direction
u(x,t) = f(tx/c) travelling along x axis
= A()exp[-i (t-s•x)] = A()exp[-i(t-k•x)]
where k= s = (/c)s is the wave number ^

18. s x sin = vt, t/x = sin/v = u sin = p
u = slowness, p = ray parameter (apparent/horizontal slowness)
rays are perpendicular to wavefronts x  s
wavefront at t+t
wavefront at t

19. p = u1 sin 1 = u2 sin 2
u = slowness, p = ray parameter (apparent/horizontal slowness) 1 v1 2 v2

20. p = u1sin 1= u2sin 2= u3sin 3
Fermat’s principle: travel time between 2 points is stationary (almost always minimum) 1 v1 2 v2 3 v3

21. Continuous Velocity Gradients
p = u0sin 0= u sin  = constant along a single ray path X v z 0  
=90o, u=utp T
dT/dX = p = ray parameter X

22. X(p) generally increases as p decreases -> dX/dp < 0v z 
p decreasing 
=90o, u=utp T
Prograde traveltime curve X

23. X(p) generally increases as p decreases but not alwaysv X z
Prograde
Retrograde T
caustics
Prograde X

24. Reduced Velocity
Prograde
Retrograde T
caustics
Prograde X
T-X/Vr X

25. X(p) generally increases as p decreases -> dX/dp < 0
Shadow
zone v X z
lvz T  X p

30. Traveltime tomographyj
Traveltime tomography
T = ∫ 1/v(s)ds = ∫u(s)ds
Tj = ∑ Gij ui
Tj = ∑ Gij ui
i=1
d=Gm
GTd=GTGm
mg=(GTG)-1GTd
j-th ray
i=1

31. Earthquake location uncertainty
2 = ∑ [ti-tip]2/i2
i expected standard deviation
2 (mbest)= ∑ [ti-tip(mbest)]2/ndf
mbest is best-fitting station
2(m) = ∑ [ti-tip]2/2 - contour! n
i=1

 n
i=1

32. Fast location:
S-P times: D ~ 8 x S-P(s)

33. Other sources of error:
Lateral velocity variations
slow fast
Station distribution

34. Emean = 1/2 A2 2
(higher frequencies carry more E!)
A2/A1= (1c1/2c2)1/2
ds2
ds1

35. 1cos1-2cos2
S’S’’=
1cos1+2cos2
2 1cos1
S’S’ =
since ucoscos
For vertical incidence (
1 - 2
A1’’=
1 + 2
2 1
A2’ =
1- 1
S’S’’vert= S’S’vert=
1+ 1+2

36. S waves vertical incidence
P waves vertical incidence:
1- 1
P’P’’vert= P’P’vert=
1+ 1+2
1- 1
S’S’’vert= S’S’vert=
1+ 1+2

37. E1flux = 1/2 c1 A12 2 cos1
E2flux = 1/2 c2 A22 2 cos2
Anorm = [E2flux/E1flux]1/2 = A2/A1 [c2cos2/ c1cos1]1/2
= Araw [c2cos2/ c1cos1]1/2
1cos1-2cos2
S’S’’norm= = S’S’’raw
1cos+2cos
2 1cos (c2cos2)1/2
S’S’norm = x
1cos1+2cos (c1cos1)1/2

38. 2 1cos
S’S’ =
1cos+2cos
What happens beyond c ? There is no transmitted
wave, and cos = (1-p2c2)1/2 becomes imaginary.
No energy is transmitted to the underlying layer,
we have total internal reflection. The vertical slowness
=(u2-p2)1/2 becomes imaginary as well. Waves with
Imaginary vertical slowness are called inhomogeneous
or evanescent waves.

39. Phase changes:
Vertical incidence, free surface:
S waves - no change in polarity
P waves - polarity change of 
Vertical incidence, impedance increases:
S waves - opposite polarity
P waves - no change in polarity
Fig 6.4
Phase advance of /2 - Hilbert Transform

40. Attenuation: scattering and intrinsic attenuation
Scattering: amplitudes reduced by scattering off small-scale
objects, integrated energy remains constant
Intrinsic:
1/Q() = -E/2E
E is the peak strain energy,
-e is energy loss per cycle
Q is the Quality factor
A(x)=A0exp(-x/2cQ)
X is distance along propagation distance
C is velocity

41. Ray methods:
t* = ∫dt/Q ( r ), A()=A0()exp(-t*/2)
i.e., higher frequencies are attenuated more!
pulse broadening