Seismology – Lecture 2 Normal modes and surface waves Barbara Romanowicz Univ. of California,...
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Transcript of Seismology – Lecture 2 Normal modes and surface waves Barbara Romanowicz Univ. of California,...
Seismology – Lecture 2Normal modes and surface waves
Barbara RomanowiczUniv. of California, Berkeley
CIDER Summer 2010 - KITP
From Stein and Wysession, 2003CIDER Summer 2010 - KITP
P S SS
Surface waves
Loma Prieta (CA) 1989 M 7 earthquake observed at KEV, Finland
From Stein and Wysession, 2003
Shallow earthquake
CIDER Summer 2010 - KITPone hour
Direction of propagation along the earth’s surface
L
Z
T
Surface waves• Arise from interaction of body waves with free
surface.• • Energy confined near the surface
• Rayleigh waves: interference between P and SV waves – exist because of free surface
• Love waves: interference of multiple S reflections. Require increase of velocity with depth
• Surface waves are dispersive: velocity depends on frequency (group and phase velocity)
• Most of the long period energy (>30 s) radiated from earthquakes propagates as surface waves
CIDER Summer 2010 - KITP
After Park et al, 2005After Park et al, 2005CIDER Summer 2010 - KITP
Free oscillations
CIDER Summer 2010 - KITP
CIDER Summer 2010 - KITP
The k’th free oscillation satisfies:
SNREI model; Solutions of the form
k = (l,m,n)
fLt
)(2
2
0 uu
0)( 20 kkk uuL
tik
keru ),,(u
CIDER Summer 2010 - KITP
Free Oscillations (Standing Waves)
€
−0ω2u = L(u)
In the frequency domain:
Free Oscillations
In a Spherical, Non-Rotating, Elastic and Isotropic Earth model,the k’th free oscillation can be described as:
l = angular order; m = azimuthal order; n = radial orderk = (l,m,n) “singlet” Degeneracy:(l,n): “multiplet” = 2l+1 “singlets ” with the same eigenfrequency nl
tik
keru ),,(u
€
uk (r,θ ,φ) =ˆ r nU l (r)Ylm (θ ,φ) +n Vl (r)∇1Yl
m (θ ,φ) −n W l (r)ˆ r ×∇1Ylm (θ ,φ)
€
k =n ω l
€
−l ≤ m ≤ l
€
Ylm (θ ,φ) = X l
m (θ )e imφ
Spheroidal modes : Vertical & Radial component
Toroidal modes : Transverse component
n T l
l : angular order, horizontal nodal planes
n : overtone number, vertical nodes
n=0n=1
CIDER Summer 2010 - KITP
Fundamentalmode
overtones
Spheroidal modes
n=0
nSl
Spatial shapes:
Depth sensitivity kernels of earth’s normal modes
53.9’
44.2’
20.9’ r=0.05m
0T22S1
0S30S2
0T4
1S2
0S5
0S0
0S43S1
2S2
1S3
0T3
Sumatra Andaman earthquake 12/26/04 M 9.3
• Rotation, ellipticity, 3D heterogeneity removes the degeneracy:
– -> For each (n, l) there are 2l+1 singlets with different frequencies
0S2 0S3
2l+1=5 2l+1=7
mode 0S3 7 singlets
Geographical sensitivity kernel K0()
0S45
0S3
ωo
Δω
frequency
Frequency shift depends only on the average structure along the vertical planecontaining the source and the receiver weighted by the depth sensitivity of the mode considered:
Mode frequency shifts
SNREI->
€
ˆ ω k ≈1
2πδω(s)ds∫
δω(θ ,φ) = Mkk (r)δm0
a
∫ (r,θ ,φ)r2dr
S
R
P(θ,Φ)
Masters et al., 1982
Anomalous splitting of core sensitive modes
Data
Model
Mantle mode
Core mode
Seismograms by mode summation
Mode Completeness:
€
u = Re{ akk
∑ (t)uk (r,θ ,ϕ )e iω k t e−α k t}
Orthonormality (L is an adjoint operator):
€
0uk'* ⋅ ukdV = δ kk '
V
∫
fLt
)(2
2
0 uu
* Denotes complex conjugate
Depends on source excitation f
Normal mode summation – 1D
A : excitationw : eigen-frequencyQ : Quality factor ( attenuation )
CIDER Summer 2010 - KITP
Spheroidal modes : Vertical & Radial component
Toroidal modes : Transverse component
n T l
l : angular order, horizontal nodal planes
n : overtone number, vertical nodes
n=0n=1
CIDER Summer 2010 - KITP
CIDER Summer 2010 - KITP
P S SS
Surface waves
Loma Prieta (CA) 1989 M 7 earthquake observed at KEV, Finland
€
u(t) = Re{ Akk
∑ e iω k t e−α k t}
Standing waves and travelling waves
Ak ---- linear combination of moment tensor elements and spherical harmonics Yl
m
When l is large (short wavelengths):
€
Ylm (θ ,ϕ ) ≈
1
π sinΔcos (l +
1
2)Δ −
π
4+
mπ
2
⎡ ⎣ ⎢
⎤ ⎦ ⎥e
imϕ
Replace x=a Δ, where Δ is angular distance and x linear distance along the earth’ssurface
Jeans’ formula : ka = l + 1/2
€
Ylm (θ ,ϕ ) ≈
1
π sin Δcos kx −
π
4+
mπ
2
⎡ ⎣ ⎢
⎤ ⎦ ⎥e
imϕ
≈1
2π sinΔe
i(kx −π
4+
mπ
2)+ e
−i(kx −π
4+
mπ
2) ⎡
⎣ ⎢
⎤
⎦ ⎥
Hence:
€
u(t) = Re{ Akk
∑ e iω k t e−α k t}
⏐ → ⏐ ∝ e i(ω k t −kx )
⏐ → ⏐ e i(ω k t +kx )
Plane wavespropagatingin opposite directions
-> Replace discrete sum over l by continuous sum over frequency (Poisson’s formula):
€
u(x, t) = S(ω)e i(ωt −kx )∫ dω
With k=k(ω) (dispersion)
€
k = k(ω)
Phase velocity:
€
C(ω) =ω
k
S is slowly varying with ω ; The main contribution to the integral is when the phase is stationary:
S is slowly varying with ω ; The main contribution to the integral is when the phase is stationary:
€
dΦ
dω= t −
dk
dωx = 0 For some frequency ωs
The energy associated with a particular group centered on ωs travels with the group velocity:
€
U(ω) =x
t=
dω
dk
Rayleigh phase velocity maps
Reference: G. Masters – CIDER 2008
Period = 50 s Period = 100 s
Group velocity maps
Period = 100 sPeriod = 50 s
Reference: G. Masters CIDER 2008
Importance of overtones for constraining structurein the transition zone
n=0: fundamental mode
n=1n=2
overtones
Overtones By including overtones, we can see into the transition zone and the top of the lower mantle.
from Ritsema et al, 2004
Ritsema et al.,2004
FundamentalModeSurfacewaves
Overtone surface waves
Body waves
120 km
325 km
600 km
1100 km
1600 km
2100 km
2800 km
Anisotropy
• In general elastic properties of a material vary with orientation
• Anisotropy causes seismic waves to propagate at different speeds– in different directions– If they have different polarizations
Types of anisotropy
• General anisotropic model: 21 independent elements of the elastic tensor cijkl
• Long period waveforms sensitive to a subset (13) of which only a small number can be resolved
– Radial anisotropy– Azimuthal anisotropy
CIDER Summer 2010 - KITP
Montagner andNataf, 1986
RadialAnisotropy
Radial (polarization) Anisotropy
• “Love/Rayleigh wave discrepancy”– Vertical axis of symmetry
• A= Vph2,
• C= Vpv2,
• F,
• L= Vsv2,
• N= Vsh2 (Love, 1911)
– Long period S waveforms can only resolve• L , N
• => = (Vsh/Vsv) 2
ln =2(ln Vsh – lnVsv)
Azimuthal anisotropy
• Horizontal axis of symmetry• Described in terms of , azimuth with
respect to the symmetry axis in the horizontal plane– 6 Terms in 2 (B,G,H) and 2 terms in 4 (E)
• Cos 2 -> Bc,Gc, Hc• Sin 2 -> Bs,Gs, Hs• Cos 4-> Ec• Sin 4 -> Es
– In general, long period waveforms can resolve Gc and Gs
Montagner and Anderson, 1989
• Vectorial tomography: – Combination radial/azimuthal (Montagner
and Nataf, 1986): – Radial anisotropy with arbitrary axis
orientation (cf olivine crystals oriented in “flow”) – orthotropic medium
– L,N, ,
x
y
z
Axis of symmetry
CIDER Summer 2010 - KITP
Montagner, 2002
= (Vsh/Vsv)2
RadialAnisotropy
Isotropic velocity
Azimuthal anisotropy
Depth= 100 km
Montagner, 2002
Ekstrom and Dziewonski, 1997
Pacific ocean radial anisotropy: Vsh > Vsv
Gung et al., 2003
Marone and Romanowicz, 2007
Absolute Plate Motion
Continuous lines: % Fo (Mg) fromGriffin et al. 2004Grey: Fo%93black: Fo%92
Yuan and Romanowicz, in press
Layer 1 thickness
Mid-continental rift zone
Trans HudsonOrogen
“Finite frequency” effects
CIDER Summer 2010 - KITP
Structure sensitivity kernels: path average approximation (PAVA)versus Finite Frequency (“Born”) kernels
SR
M
SR
M
PAVA
2DPhasekernels
Panning et al., 2009
Waveform tomography
observed
synthetic
Waveform Tomography