Lecture 13

Seismology

　

‘Earthquakes don’t kill people, structures do’ 　　

1999 Izmit, Turkey earthquake

The number one cause of damage from earthquakes is due to collapse of buildings from ground shaking (The number two cause is tsunamis)

Credit: Walter Mooney, USGS

Wave propagation

• When a wavefront hits the interface

between 2 medias with different

physical properties (⇒ different

seismic velocities)

– Some of the energy is reflected back in

medium 1 = reflection

– Some of the energy goes trough the

interface into medium 2 = refraction

• Reflection: iincident = ireflection

• Refraction:

– Snell-Descartes law:

– If V2>V1 (usual case in the Earth), then

r > i!

sini

V1

=sin r

V2

Incident ray

refle

cted

ray

refracted ray

interface

velocity 1

velocity 2

(v2 > v1)

ii

r

Wave propagation

• When a wavefront hits the interface

between 2 medias with different

physical properties (⇒ different

seismic velocities)

– Some of the energy is reflected back in

medium 1 = reflection

– Some of the energy goes trough the

interface into medium 2 = refraction

• Reflection: iincident = ireflection

• Refraction:

– Snell-Descartes law:

– If V2>V1 (usual case in the Earth), then

r > i!

sini

V1

=sin r

V2

Incident ray

refle

cted

ray

refracted ray

interface

velocity 1

velocity 2

(v2 > v1)

ii

r

If V2 > V1, r > i If V2 < V1, r < i When r = 90o, Head waves

Waves Through the Earth

Reflection & Refraction Snell’s law:

Wave

propagation

• Seismic source at O

• Homogeneous and uniformly thick layer with

velocity V1 on top of layer V2 (V2 >V1).

• Let’s consider seismic rays with increasing

incidence angle i:

– Between O and C:

• Simple reflection and refraction

• r increases as i increases (because V2>V1)

– At C and after:

• i is such that r = π/2

• i0 = critical angle of incidence

• Refracted ray travels at 90o from normal to interface

=> parallel to interface (but in lower layer) = critically

refracted rays

• Refracted rays accompanied by reflections upward

with angle i0 = head wave

• Still reflections = supercritical reflections (or wide-

angle reflections)

critically refracted ray

subcritical reflection

incident

rays

reflected

rays

refracted rays

Osupercritical reflection

“head waves”

V1

V2

critical

distance

C

io

ic

sin(r) = (V2/V1) x sin(i) with V1=4.0 km/s V2=5.7 km/s

0

10

20

30

40

50

60

70

80

90

0 10 20 30 40 50 60 70 80 90

incident angle (i)

refr

action a

ngle

(r)

critical angle i0=53.13o

Credit: Lidunka Vočadlo, UCL

Wave

propagation

• Seismic source at O

• Homogeneous and uniformly thick layer with

velocity V1 on top of layer V2 (V2 >V1).

• Let’s consider seismic rays with increasing

incidence angle i:

– Between O and C:

• Simple reflection and refraction

• r increases as i increases (because V2>V1)

– At C and after:

• i is such that r = π/2

• i0 = critical angle of incidence

• Refracted ray travels at 90o from normal to interface

=> parallel to interface (but in lower layer) = critically

refracted rays

• Refracted rays accompanied by reflections upward

with angle i0 = head wave

• Still reflections = supercritical reflections (or wide-

angle reflections)

critically refracted ray

subcritical reflection

incident

rays

reflected

rays

refracted rays

Osupercritical reflection

“head waves”

V1

V2

critical

distance

C

io

ic

sin(r) = (V2/V1) x sin(i) with V1=4.0 km/s V2=5.7 km/s

0

10

20

30

40

50

60

70

80

90

0 10 20 30 40 50 60 70 80 90

incident angle (i)

refr

action a

ngle

(r)

critical angle i0=53.13o

Refraction• Homogeneous medium of uniform velocity V1 over

medium V2: travel time of direct, refracted and reflectedrays?

• t as a function of distance x along the profile = t-x curve

• Direct ray:

⇒ t-x curve: straight line, slope 1/V1

• Refracted ray (head wave):

⇒ t-x curve: straight line, slope 1/V2⇒ Intercept time tint=> h1

!

t =x

V1

!

t =SC

V1

+CD

V2

+D " D

V1

= 2h1

V1cosi

c

+CD

V2

with CD = x # 2h1tani

c

$ t = 2h1

V1cosi

c

+x

V2

#2h

1tani

c

V2

sinceV1

V2

= sinic

$ t =x

V2

+2h

1cosi

c

V1

=x

V2

+2h

1V2

2#V

1

2

V1V2

=x

V2

+ tint

travel time t

xO

refracted rays

slope = 1/V2

direct arrivals

slope = 1/V1

tint

critical

distance

cross-over

distance

direct rayS D’

h1

refr

acte

d ra

y

O x

ic ic

C D

V1

V2

V12t2 = 4h1

2 + x2

Reflection

• Horizontal reflector

• Homogeneous medium of uniform velocity V

• Two-way travel time the source and receiver:

direct ray

reflector

source receiver

hD

refl

ecte

d ra

y

!

t = 2D

Vand D = h

2+

x

2

"

# $ %

& '

2

( t =2

Vh2

+x

2

"

# $ %

& '

2

=2h

V1+

x2

4h2

= to1+

x2

4h2

O x

• t-x curve:

– At x=0, vertical reflection: t = to= 2h / V

– When x increases, normal move out

• For reflected arrivals, the t-x equation can also be written:

⇒ Hyperbola, asymptotic to direct arrivals

!

t2

to

2"x2

4h2

=1

x

travel time t

xO

direct arrivals

slope = 1/V1t o

= 2

h/V

reflection

hyperbola

N(t) =C1

t

N(t) =C1

(C2 + t)p

t =2

V1

rh

21 +

x

2

4

2

Direct wave:

Reflected wave:

Head wave:

N(t) =C1

t

N(t) =C1

(C2 + t)p

t =2

V1

rh

21 +

x

2

4

t =x

V2+

2h1

qV

22 �V

21

V1V2

2

1/V2

Refraction• Homogeneous medium of uniform velocity V1 over

medium V2: travel time of direct, refracted and reflectedrays?

• t as a function of distance x along the profile = t-x curve

• Direct ray:

⇒ t-x curve: straight line, slope 1/V1

• Refracted ray (head wave):

⇒ t-x curve: straight line, slope 1/V2⇒ Intercept time tint=> h1

!

t =x

V1

!

t =SC

V1

+CD

V2

+D " D

V1

= 2h1

V1cosi

c

+CD

V2

with CD = x # 2h1tani

c

$ t = 2h1

V1cosi

c

+x

V2

#2h

1tani

c

V2

sinceV1

V2

= sinic

$ t =x

V2

+2h

1cosi

c

V1

=x

V2

+2h

1V2

2#V

1

2

V1V2

=x

V2

+ tint

travel time t

xO

refracted rays

slope = 1/V2

direct arrivals

slope = 1/V1

tint

critical

distance

cross-over

distance

direct rayS D’

h1

refr

acte

d ra

y

O x

ic ic

C D

V1

V2O

N N’

Tracing rays through the Earth

• Ray parameter p:

sin i1 = sin i2= …= sin in = const. = p

V1 V2 Vn

⇒ p is constant for a given ray

• If V increases continuouslywith depth ⇒ the ray flattens

and eventually bends upwardsuntil critical refraction isreached

i1

i2

i3

i4

icritical

critical ref raction

source receiv er

V1

V2

V3

V4

source receiv er

Velocity increases

continuously with depthTracing rays through the Earth

• Ray parameter p:

sin i1 = sin i2= …= sin in = const. = p

V1 V2 Vn

⇒ p is constant for a given ray

• If V increases continuouslywith depth ⇒ the ray flattens

and eventually bends upwardsuntil critical refraction isreached

i1

i2

i3

i4

icritical

critical ref raction

source receiv er

V1

V2

V3

V4

source receiv er

Velocity increases

continuously with depth

Ray parameter (p):

Jeffreys-Bullen traveltime plot

Credit: Lidunka Vočadlo, UCL

Wave propagation in the Earth

• A classic black and white picture….

P a P-wave in the mantle

S an S-wave in the mantle

K a P-wave through the outer core

I a P-wave through the inner core

J an S-wave through the inner core

c a reflection from the CMB

i a reflection from the OCB-ICB

Mode conversion

Heterogeneities in D00 may explain part of the traveltimes anomalies(Breger et al., 2000a,b; Romanowicz and Breger, 2000), but differ-ential traveltimes of PKPDF and PKPBC waves have provided compel-ling evidences that the anomalies indeed originate from the innercore (Shearer et al., 1988; Creager, 1992; Song and Helmberger,1993). Shear-waves traveling in the inner core are extremely difficultto observe, but there is one report of observation of shear-wavesplitting in the inner core, which suggests an average shear-waveanisotropy of 1% in the western hemisphere (Wookey and Helffrich,2008). The inner core also exhibits a clear attenuation anisotropy,P-waves traveling parallel to the Earth’s rotation axis being moreattenuated (Souriau and Romanowicz, 1996, 1997; Cormier et al.,1998; Oreshin and Vinnik, 2004; Yu and Wen, 2006). The positivecorrelation between velocity and attenuation is the opposite of whatis observed in the mantle, thus suggesting a different attenuationmechanism.

There is a general agreement about weaker anisotropy in theeastern hemisphere than in the western hemisphere (Tanaka andHamaguchi, 1997; Creager, 1999; Garcia and Souriau, 2000; Niuand Wen, 2001; Deuss et al., 2010; Leykam et al., 2010; Irving andDeuss, 2011), the two hemispheres being defined in reference tothe Greenwich meridian. The uppermost inner core is isotropic orweakly anisotropic and highly attenuating (Shearer, 1994; Songand Helmberger, 1995b; Creager, 1999; Garcia and Souriau,2000), with again a hemispherical structure, the eastern hemi-sphere being faster and more attenuating. The hemisphericalpattern of traveltime anomalies can be explained by either adifferent anisotropy level (Creager, 1999), or by a hemisphericalvariation of the thickness of the isotropic layer, with a thicknessof ! 100 km and ! 400 km in the eastern and western hemi-spheres, respectively (Creager, 2000; Garcia and Souriau, 2000).

The depth dependence of anisotropy in the deeper inner core ismore controversial. There are evidences for the presence of aninnermost inner core of radius 300–600 km exhibiting a possiblydifferent anisotropy geometry (Ishii and Dziewonski, 2002, 2003;Beghein and Trampert, 2003; Sun and Song, 2008a,b; Niu andChen, 2008) and a lower attenuation (Li and Cormier, 2002;Cormier and Li, 2002; Cormier and Stroujkova, 2005).

The North–South orientation of the anisotropy suggests thateither Earth’s rotation or the magnetic field plays a role in thedevelopment of inner core anisotropy, which in turn suggestspossible couplings with the outer core and the dynamo. A numberof models have been proposed so far (Sections 3–6), but we arestill short of a model able of explaining all the complexities of theinner core structure in a self-consistent way.

2.2. Is there a stably stratified layer at the base of the outer core?If yes, how has it been formed and sustained?

Model PREM (Dziewonski and Anderson, 1981) exhibits onlyvery slight deviations from a well-mixed adiabatic state in theouter core. This is consistent with the idea of vigorous coreconvection, which is expected to allow only minute deviationsfrom an adiabatic profile (e.g. Stevenson, 1987) except in verythin boundary layers at the top and bottom of the core. YetSouriau and Poupinet (1991) have found evidences of significantdeviations from PREM of the P-wave velocity profile at the base ofthe outer core. Traveltimes of PKPBC and PKPCdiff (P-wavesdiffracted along the ICB) suggest the presence of a region of lowP-wave velocity gradient in the bottom ! 150 km of the outercore. This observation has then been confirmed by a number ofstudies (Song and Helmberger, 1992; Yu et al., 2005; Zou et al.,2008; Cormier, 2009; Cormier et al., 2011), and has been includedin global models AK 135 (Kennett et al., 1995) and PREM2 (Songand Helmberger, 1995a). Yu et al. (2005) suggested that this layerexhibits an hemispherical structure, with the P-wave velocityprofile being closer to PREM in the eastern hemisphere. This isproblematic from a geodynamical point of view, because of thedifficulty to sustain large lateral density variations in the outercore, unless the viscosity in this region is significantly larger thanusually thought (Cormier, 2009). Cormier et al. (2011) demon-strated however that the hemispherical structure proposed by Yuet al. (2005) is not required by the data.

The observation of this anomalous layer, sometimes called the‘F-layer’,1 is puzzling, and seems very difficult to reconcile withour current view of outer core convection, where crystallization ofthe inner core is thought to drive convection through the releaseof buoyant, light element rich liquid. The low P-wave velocitygradient is most probably indicative of a stable density stratifica-tion, which is just the opposite of what we expect. The layer isalso much thicker than what we can expect for a convectivethermo-chemical boundary layer. Models for the formation of thislayer are discussed in Section 7.

2.3. Is the inner core differentially rotating? What does it imply?

Following theoretical (Gubbins, 1981) and numerical (Glatzmaierand Roberts, 1996) predictions of inner core differential rotation withrespect to the mantle, a number of seismological observations have

A

B

C

D

F

1100

1150

1200

1250

1300

1350

Tra

vel t

ime

(s)

120 140 160 180

0˚

30˚

60˚90˚

120˚

150˚

180˚

0˚

30˚

60˚90˚

120˚

150˚

180˚

0˚

30˚

60˚90˚

120˚

150˚

180˚Distance ∆ (°)

PKP A

BPKP Cdif

f

PKIKP

PKiKP

PKP BC

Fig. 1. (a) Traveltimes of the main core phases as a function of epicentral distanceD. (b)–(d) Ray paths of the main core phases with epicentral distance equal,respectively, to 1401, 1501, and 1701. In the epicentral distance range 145–1551,two different PKP phases coexist: PKPAB (red) samples mostly the outer part of theouter core, while PKPBC (orange) samples the base of the outer core. Wheneverpossible, PKIKP-waves (also called PKPDF) are compared to a reference PKP phase,either PKPAB or PKPBC, to minimize the effect of source mislocation and mantleheterogeneities. This is possible for only a limited range of epicentral distances (orequivalently, of the depth of the ray turning point). The PKPBC phase exists only forepicentral distances between 1451 and 1551, which correspond to a turning pointo400 km below the ICB. For deep ray paths (epicentral distance 41551, turningpoint 4400 km below the ICB), PKIKP and the reference phase PKPAB sample quitedifferent regions in D00 , as can be seen in (d). PKPCdiff is a compressional wavediffracted around the inner core boundary. (For interpretation of the references tocolor in this figure caption, the reader is referred to the web version of this article.)

1 In reference to a layer present in early seismological models between theouter core and the inner core, the E and G layers in Bullen’s nomenclature (Jeffreys,1939; Bullen, 1947). The original F-layer had much more dramatic features,including a strong P-wave velocity discontinuity at its boundary with the outercore. The PKiKP precursors interpreted as a discontinuity at the top of the F-layerwere later shown to be due to scattering of PKP-waves in D00 (Cleary and Haddon,1972; King et al., 1973), and the F-layer disappeared from global Earth models.

R. Deguen / Earth and Planetary Science Letters 333–334 (2012) 211–225212

(http://www.ucl.ac.uk)

(Deguen, EPSL, 2012)

ever-greater distances up to 180!. There are at least twobranches of the t–" curve for "#143!, corresponding tothe PKP and PKIKP phases, respectively (Fig. 3.77). Infact, depending how the transitional region F is modelled,the t–" curve near 143! can have several branches.

The edges of the shadow zone defined by P and PKPphases are not sharp. One reason is the intrusion ofPKIKP phases at the 143! edge. Another is the effect ofdiffraction of P-waves at the 103! edge (Fig. 3.79). Thebending of plane waves at an edge into the shadow of anobstruction was described in Section 3.6.2.3, andexplained with the aid of Huygen’s principle. Thediffraction of plane waves is called Fraunhofer diffraction.When their source is not at infinity, waves must behandled as spherical waves. Spherical wavefronts thatpass an obstacle are also diffracted. This type of behavioris called Fresnel diffraction, and it is also explainable withHuygen’s principle as the product of interference betweenthe primary wavefront and secondary waves generated atthe obstacle. Wave energy penetrates into the shadowof the obstacle, as though the wavefront were bent aroundthe edge. In this way very deep P-waves are diffractedaround the core and into the shadow zone. The intensityof the diffracted rays falls off with increasing angular dis-tance from the diffracting edge, in this case the core–mantle boundary. Modern instrumentation enablesdetection of long-period diffracted P-waves to large epi-central distances (Figs. 3.77, 3.79). The velocity structureabove the core–mantle boundary, in particular in the D$-layer (Section 3.7.5.3), has a strong influence on the raypaths, travel-times and waveforms of the diffracted waves.

3.7.3 Radial variations of seismic velocities

Models of the radial variations of physical parametersinside the Earth implicitly assume spherical symmetry.They are therefore “average” models of the Earth that do

not take into account lateral variations (e.g., of velocityor density) at the same depth. This is a necessary first stepin approaching the true distributions as long as lateralvariations are relatively small. This appears to be the case;although geophysically significant, the lateral variationsin physical properties remain within a few percent of theaverage value at any depth.

There are two main ways to determine the distribu-tions of body-wave velocities in a spherically symmetricEarth. They are referred to as forward and inverse model-ling. Both methods have to employ the same sets of obser-vations, which are the travel-times of different seismicphases to known epicentral distances. The forward tech-nique starts with a known or assumed variation ofseismic velocities and calculates the corresponding travel-times. The inversion method starts with the observed t–"curves and computes a model of the velocity distributionsthat could produce the curves. The inversion method isthe older one, in use since the early part of the twentiethcentury, and forms an important branch of mathematicaltheory. Forward modelling is a more recent method thathas been successfully employed since the advent of pow-erful computers.

3.7.3.1 Inversion of travel-time versus distance curves

In 1907 the German geophysicist E. Wiechert, buildingupon an evaluation of the Benndorf problem (Eq.(3.147)) by the mathematician G. Herglotz, developedan analytical method for computing the internal distrib-utions of seismic velocities from observations made atthe Earth’s surface. The technique is called inversion oftravel-times, and it is considered one of the classicalmethods of geophysics. The observational data consistof the t–" curves for important seismic phases(Fig. 3.77). The clues to deciphering the velocity dis-tributions were the Benndorf relationship for the ray

190 Seismology and the internal structure of the Earth

20°

40°

60°

80° 100°110°

120°

130°

140°

150°

160°

170°

180°

E

P

M

D

C

B

A

O

N

H

G

L

K

JI

F

LKJI

DC

BA

ON

HE GF

P

M

diffracted P

mantle E F Gcore

outer inner

Fig. 3.79 The wave paths ofsome P, PKP, and PKIKP rays(after Gutenberg, 1959).

Lowrie, Fundamentals of Geophysics

Heterogeneities in D00 may explain part of the traveltimes anomalies(Breger et al., 2000a,b; Romanowicz and Breger, 2000), but differ-ential traveltimes of PKPDF and PKPBC waves have provided compel-ling evidences that the anomalies indeed originate from the innercore (Shearer et al., 1988; Creager, 1992; Song and Helmberger,1993). Shear-waves traveling in the inner core are extremely difficultto observe, but there is one report of observation of shear-wavesplitting in the inner core, which suggests an average shear-waveanisotropy of 1% in the western hemisphere (Wookey and Helffrich,2008). The inner core also exhibits a clear attenuation anisotropy,P-waves traveling parallel to the Earth’s rotation axis being moreattenuated (Souriau and Romanowicz, 1996, 1997; Cormier et al.,1998; Oreshin and Vinnik, 2004; Yu and Wen, 2006). The positivecorrelation between velocity and attenuation is the opposite of whatis observed in the mantle, thus suggesting a different attenuationmechanism.

There is a general agreement about weaker anisotropy in theeastern hemisphere than in the western hemisphere (Tanaka andHamaguchi, 1997; Creager, 1999; Garcia and Souriau, 2000; Niuand Wen, 2001; Deuss et al., 2010; Leykam et al., 2010; Irving andDeuss, 2011), the two hemispheres being defined in reference tothe Greenwich meridian. The uppermost inner core is isotropic orweakly anisotropic and highly attenuating (Shearer, 1994; Songand Helmberger, 1995b; Creager, 1999; Garcia and Souriau,2000), with again a hemispherical structure, the eastern hemi-sphere being faster and more attenuating. The hemisphericalpattern of traveltime anomalies can be explained by either adifferent anisotropy level (Creager, 1999), or by a hemisphericalvariation of the thickness of the isotropic layer, with a thicknessof ! 100 km and ! 400 km in the eastern and western hemi-spheres, respectively (Creager, 2000; Garcia and Souriau, 2000).

The depth dependence of anisotropy in the deeper inner core ismore controversial. There are evidences for the presence of aninnermost inner core of radius 300–600 km exhibiting a possiblydifferent anisotropy geometry (Ishii and Dziewonski, 2002, 2003;Beghein and Trampert, 2003; Sun and Song, 2008a,b; Niu andChen, 2008) and a lower attenuation (Li and Cormier, 2002;Cormier and Li, 2002; Cormier and Stroujkova, 2005).

The North–South orientation of the anisotropy suggests thateither Earth’s rotation or the magnetic field plays a role in thedevelopment of inner core anisotropy, which in turn suggestspossible couplings with the outer core and the dynamo. A numberof models have been proposed so far (Sections 3–6), but we arestill short of a model able of explaining all the complexities of theinner core structure in a self-consistent way.

2.2. Is there a stably stratified layer at the base of the outer core?If yes, how has it been formed and sustained?

Model PREM (Dziewonski and Anderson, 1981) exhibits onlyvery slight deviations from a well-mixed adiabatic state in theouter core. This is consistent with the idea of vigorous coreconvection, which is expected to allow only minute deviationsfrom an adiabatic profile (e.g. Stevenson, 1987) except in verythin boundary layers at the top and bottom of the core. YetSouriau and Poupinet (1991) have found evidences of significantdeviations from PREM of the P-wave velocity profile at the base ofthe outer core. Traveltimes of PKPBC and PKPCdiff (P-wavesdiffracted along the ICB) suggest the presence of a region of lowP-wave velocity gradient in the bottom ! 150 km of the outercore. This observation has then been confirmed by a number ofstudies (Song and Helmberger, 1992; Yu et al., 2005; Zou et al.,2008; Cormier, 2009; Cormier et al., 2011), and has been includedin global models AK 135 (Kennett et al., 1995) and PREM2 (Songand Helmberger, 1995a). Yu et al. (2005) suggested that this layerexhibits an hemispherical structure, with the P-wave velocityprofile being closer to PREM in the eastern hemisphere. This isproblematic from a geodynamical point of view, because of thedifficulty to sustain large lateral density variations in the outercore, unless the viscosity in this region is significantly larger thanusually thought (Cormier, 2009). Cormier et al. (2011) demon-strated however that the hemispherical structure proposed by Yuet al. (2005) is not required by the data.

The observation of this anomalous layer, sometimes called the‘F-layer’,1 is puzzling, and seems very difficult to reconcile withour current view of outer core convection, where crystallization ofthe inner core is thought to drive convection through the releaseof buoyant, light element rich liquid. The low P-wave velocitygradient is most probably indicative of a stable density stratifica-tion, which is just the opposite of what we expect. The layer isalso much thicker than what we can expect for a convectivethermo-chemical boundary layer. Models for the formation of thislayer are discussed in Section 7.

2.3. Is the inner core differentially rotating? What does it imply?

Following theoretical (Gubbins, 1981) and numerical (Glatzmaierand Roberts, 1996) predictions of inner core differential rotation withrespect to the mantle, a number of seismological observations have

A

B

C

D

F

1100

1150

1200

1250

1300

1350

Tra

vel t

ime

(s)

120 140 160 180

0˚

30˚

60˚90˚

120˚

150˚

180˚

0˚

30˚

60˚90˚

120˚

150˚

180˚

0˚

30˚

60˚90˚

120˚

150˚

180˚Distance ∆ (°)

PKP A

BPKP Cdif

f

PKIKP

PKiKP

PKP BC

Fig. 1. (a) Traveltimes of the main core phases as a function of epicentral distanceD. (b)–(d) Ray paths of the main core phases with epicentral distance equal,respectively, to 1401, 1501, and 1701. In the epicentral distance range 145–1551,two different PKP phases coexist: PKPAB (red) samples mostly the outer part of theouter core, while PKPBC (orange) samples the base of the outer core. Wheneverpossible, PKIKP-waves (also called PKPDF) are compared to a reference PKP phase,either PKPAB or PKPBC, to minimize the effect of source mislocation and mantleheterogeneities. This is possible for only a limited range of epicentral distances (orequivalently, of the depth of the ray turning point). The PKPBC phase exists only forepicentral distances between 1451 and 1551, which correspond to a turning pointo400 km below the ICB. For deep ray paths (epicentral distance 41551, turningpoint 4400 km below the ICB), PKIKP and the reference phase PKPAB sample quitedifferent regions in D00 , as can be seen in (d). PKPCdiff is a compressional wavediffracted around the inner core boundary. (For interpretation of the references tocolor in this figure caption, the reader is referred to the web version of this article.)

1 In reference to a layer present in early seismological models between theouter core and the inner core, the E and G layers in Bullen’s nomenclature (Jeffreys,1939; Bullen, 1947). The original F-layer had much more dramatic features,including a strong P-wave velocity discontinuity at its boundary with the outercore. The PKiKP precursors interpreted as a discontinuity at the top of the F-layerwere later shown to be due to scattering of PKP-waves in D00 (Cleary and Haddon,1972; King et al., 1973), and the F-layer disappeared from global Earth models.

R. Deguen / Earth and Planetary Science Letters 333–334 (2012) 211–225212

Seismology and the Earth’s Deep Interior Seismogram Interpretation

Wavefields in the Earth: SH wavesWavefields in the Earth: SH waves

Red and yellowcolor denote positive and negative displacement, respectively.

Wavefield for earthquake at 600km depth.

Seismology and the Earth’s Deep Interior Seismogram Interpretation

Wavefields in the Earth: SH wavesWavefields in the Earth: SH waves

Red and yellowcolor denote positive and negative displacement, respectively.

Wavefield for earthquake at 600km depth.

Seismology and the Earth’s Deep Interior Seismogram Interpretation

Wavefields in the Earth: SH wavesWavefields in the Earth: SH waves

Red and yellowcolor denote positive and negative displacement, respectively.

Wavefield for earthquake at 600km depth.

Seismology and the Earth’s Deep Interior Seismogram Interpretation

Wavefields in the Earth: SH wavesWavefields in the Earth: SH waves

Red and yellowcolor denote positive and negative displacement, respectively.

Wavefield for earthquake at 600km depth.

Seismology and the Earth’s Deep Interior Seismogram Interpretation

Wavefields in the Earth: SH wavesWavefields in the Earth: SH waves

Red and yellowcolor denote positive and negative displacement, respectively.

Wavefield for earthquake at 600km depth.

Credit: Heiner Igel, UMunich

Seismology and the Earth’s Deep Interior Seismogram Interpretation

SH waves: seismogramsSH waves: seismograms

SH-seismograms for a source at 600km depth

Credit: Heiner Igel, UMunich

Velocity-depth structure

• Applied to the Earth:

– Major discontinuities in the seismic structure of the Earth ⇔ discontinuitiesin mineralogy/petrology

• Moho, 410 km, 660 km, D”, CMB

– IASP91 model (below), PREM (Preliminary Reference Earth Model)

– These models assume perfect spherical symmetry.Preliminary Reference Earth Model (PREM)

410 – Olivine to Spinel 660 – Spinel to Perovskite (Bridgmanite) D’’ – Perovskite to Post-perovskite (PPV)

Credit: Lidunka Vočadlo, UCL