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**Solid Earth**
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**Review article**
05 Feb 2019

**Review article** | 05 Feb 2019

Power spectra of random heterogeneities in the solid earth

^{1}Geophysics, Science, Tohoku University, Aoba-ku, Sendai-shi, Miyagi-ken, 980-8578, Japan^{*}Invited contribution by Haruo Sato, recipient of the EGU Beno Gutenberg Medal 2018.

^{1}Geophysics, Science, Tohoku University, Aoba-ku, Sendai-shi, Miyagi-ken, 980-8578, Japan^{*}Invited contribution by Haruo Sato, recipient of the EGU Beno Gutenberg Medal 2018.

**Correspondence**: Haruo Sato (haruo.sato.e2@tohoku.ac.jp)

**Correspondence**: Haruo Sato (haruo.sato.e2@tohoku.ac.jp)

Abstract

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Recent seismological observations focusing on the collapse of an impulsive
wavelet revealed the existence of small-scale random heterogeneities in the earth medium.
The radiative transfer theory (RTT) is often used for the study of the propagation and
scattering of wavelet intensities, the mean square amplitude envelopes through random
media. For the statistical characterization of the power spectral density function (PSDF)
of the random fractional fluctuation of velocity inhomogeneities in a 3-D space, we use
an isotropic von Kármán-type function characterized by three parameters: the root
mean square (RMS) fractional velocity fluctuation, the characteristic length, and the
order of the modified Bessel function of the second kind, which leads to the power-law
decay of the PSDF at wavenumbers higher than the corner. We compile reported statistical
parameters of the lithosphere and the mantle based on various types of measurements for a
wide range of wavenumbers: photo-scan data of rock samples; acoustic well-log data; and
envelope analyses of cross-hole experiment seismograms, regional seismograms, and
teleseismic waves based on the RTT. Reported exponents of wavenumber are distributed
between −3 and −4, where many of them are close to −3. Reported RMS fractional
fluctuations are on the order of 0.01–0.1 in the crust and the upper mantle. Reported
characteristic lengths distribute very widely; however, each one seems to be restricted
by the dimension of the measurement system or the sample length. In order to grasp the
spectral characteristics, eliminating strong heterogeneity data and the lower mantle
data, we have plotted all the reported PSDFs of the crust and the upper mantle against
wavenumber for a wide range (10^{−3}–10^{8} km^{−1}). We find that the spectral
envelope of those PSDFs is well approximated by the inverse cube of wavenumber. It suggests that the earth-medium randomness has a broad
spectrum. In theory, we need to re-examine the applicable range of the Born approximation
in the RTT when the wavenumber of a wavelet is much higher than the corner. In
observation, we will have to carefully measure the PSDF on both sides of the corner. We
may consider the obtained power-law decay spectral envelope as a reference for studying
the regional differences. It is interesting to study what kinds of geophysical processes
created the observed power-law spectral envelope at different scales and in different
geological environments in the solid earth medium.

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How to cite.

Sato, H.: Power spectra of random heterogeneities in the solid earth, Solid Earth, 10, 275–292, https://doi.org/10.5194/se-10-275-2019, 2019.

1 Introduction

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The first image of the solid earth is composed of spherical shells, for example,
PREM (preliminary reference Earth model) (Dziewonski and Anderson, 1981). As
seismic networks were deployed on the regional scale and worldwide, the velocity
tomography based on the ray tracing method revealed 3-D heterogeneous structures at
various scales; however, spatial variations in the resultant velocity structure are
essentially smooth compared with seismic wavelengths. Aki and Chouet (1975) first put a
focus on long-lasting coda waves of small earthquakes and interpreted them as scattered
waves by small-scale random heterogeneities. They proposed to measure the scattering
coefficient *g*, the scattering power per unit volume as a measure of medium
heterogeneity. They analyzed the mean square (MS) amplitude time trace of coda waves as
an incoherent sum of scattered wave power by using the Born approximation
(e.g., Chernov, 1960), which is a simplified version of the radiative transfer
theory (RTT). There have been many measurements of the total scattering coefficient
*g*_{iso} supposing isotropic scattering (e.g., Sato, 1977a) in various
seismotectonic environments. The total scattering coefficient of *S* waves is reported to
be on the order of 10^{−2} km^{−1} for 1–20 Hz in the lithosphere, and it marks a
higher value beneath active volcanoes (e.g., Sato et al., 2012; Yoshimoto and Jin, 2008).
There were precise measurements of regional variations in *g*_{iso} as
Carcolé and Sato (2010) in Japan and Eulenfeld and Wegler (2017) in the US.
Hock et al. (2004) analyzed medium heterogeneity in Europe from the analyses of
teleseismic waves using the modified energy flux model (Korn, 1993). There were also
measurements of the anisotropic scattering coefficient from the analysis of *S* coda
envelopes Jing et al. (e.g., 2014); Zeng (e.g., 2017).

Aki and Chouet (1975) derived the angular dependence of the scattering coefficient of
scalar waves from the power spectral density function (PSDF) of the fractional velocity
fluctuation using the Born approximation. Sato (1984) extended the envelope
synthesis of scalar waves to the whole envelope synthesis of three-component seismograms
from the *P* onset to *S* coda on the bases of the single scattering approximation of the
RTT. His syntheses explain how seismogram envelopes in different back azimuths vary
depending on the source fault mechanism. Extension to the multiple scattering case was
developed by using Stokes parameters
(e.g., Margerin et al., 2000; Margerin, 2005; Przybilla et al., 2009; Sanborn et al., 2017).
We also note that Monte Carlo simulation was developed to stochastically solve the RTT
(e.g., Hoshiba et al., 1991; Gusev and Abubakirov, 1987; Yoshimoto, 2000). For the data
processing, it is more appropriate to stack MS envelopes of observed seismograms for
comparison with the averaged intensity time traces stochastically synthesized by the RTT
(e.g., Shearer and Earle, 2004; Rost et al., 2006; da Silva et al., 2018).

When the central wavenumber of a wavelet increases much larger than the corner wavenumber
of the PSDF, the wavelet around the peak value is mostly composed of narrow-angle
scattering around the forward direction. In such a case, the Born approximation becomes
inappropriate; however, the phase shift modulation based on the parabolic approximation
is useful, which is called the phase screen approximation. As an extension of the RTT
with the phase screen approximation, the Markov approximation was also used for the
analysis of envelope broadening and peak delay with increasing travel distance
(e.g., Sato, 1989; Saito et al., 2002; Takahashi et al., 2009). Kubanza et al. (2007)
measured regional differences in the lithospheric heterogeneity from the partitioning of
seismic energy of teleseismic *P* waves into the vertical and transverse components based
on the Markov approximation.

There have been various kinds of measurements of the PSDF of the random velocity fluctuation, where the PSDF is often supposed to be a von Kármán type. In the following section, the main objective is to compile reported PSDF measurements in various scales in different geological environments of the solid earth: photo scanning of small rock samples, acoustic well logs, array analyses of teleseismic waves, waveform analyses using finite difference (FD) simulations, and analyses of seismogram envelopes on the basis of the RTT. We enumerate their statistical parameters and plot their PSDFs against wavenumber. We will show that the envelope of all the PSDFs is well approximated by a power-law decay curve. Then, we will discuss the results obtained and a few problems in the envelope synthesis theory for such random media and the geophysical origin of such power spectra.

2 Statistical characterization of random media

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We consider the propagation of scalar waves as a simple model, where the
inhomogeneous velocity is given by $V\left(\mathit{x}\right)={V}_{\mathrm{0}}(\mathrm{1}+\mathit{\xi}(\mathit{x}\left)\right)$. The fractional fluctuation
*ξ*(** x**) is supposed to be a random function of space. We imagine an
ensemble of random media {

$$\begin{array}{}\text{(1a)}& {\displaystyle}R\left(\mathit{x}\right)=R\left(r\right)=\u2329\mathit{\xi}\left(\mathit{y}\right)\mathit{\xi}(\mathit{y}+\mathit{x})\u232a,\end{array}$$

where $r=\left|\mathit{x}\right|$. The MS fractional fluctuation as a measure of the strength of randomness is supposed to be small, and ${\mathit{\epsilon}}^{\mathrm{2}}\equiv R\left(\mathrm{0}\right)\ll \mathrm{1}$. The Fourier transform of ACF gives the PSDF:

$$\begin{array}{}\text{(1b)}& {\displaystyle}P\left(\mathit{m}\right)=P\left(m\right)=\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iiint}}R\left(\mathit{x}\right){e}^{-i\mathit{m}\mathit{x}}\mathrm{d}\mathit{x},\end{array}$$

where wavenumber $m=\left|\mathit{m}\right|$.
In some literature, (2*π*)^{−3} is used as a prefactor in the right-hand side of
Eq. (2b).

There are several types of PSDF and ACF characterized by a few parameters.

The ACF is written by using a modified Bessel
function of the second kind of order *κ* and characteristic length *a*:

$$\begin{array}{}\text{(2a)}& {\displaystyle}R\left(r\right)={\displaystyle \frac{{\mathrm{2}}^{\mathrm{1}-\mathit{\kappa}}}{\mathrm{\Gamma}\left(\mathit{\kappa}\right)}}{\mathit{\epsilon}}^{\mathrm{2}}{\left({\displaystyle \frac{r}{a}}\right)}^{\mathit{\kappa}}{K}_{\mathit{\kappa}}\left({\displaystyle \frac{r}{a}}\right)\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathrm{for}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathit{\kappa}>\mathrm{0},\end{array}$$

which is an exponential type $R\left(r\right)={\mathit{\epsilon}}^{\mathrm{2}}{e}^{-r/a}$ when
$\mathit{\kappa}=\mathrm{1}/\mathrm{2}$. In the case of space dimension *d*, the PSDF is

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}P\left(m\right)={\displaystyle \frac{{\mathrm{2}}^{d}{\mathit{\pi}}^{\frac{d}{\mathrm{2}}}\mathrm{\Gamma}(\mathit{\kappa}+\frac{d}{\mathrm{2}}){\mathit{\epsilon}}^{\mathrm{2}}\phantom{\rule{0.125em}{0ex}}{a}^{d}}{\mathrm{\Gamma}\left(\mathit{\kappa}\right){\left(\mathrm{1}+{a}^{\mathrm{2}}{m}^{\mathrm{2}}\right)}^{\mathit{\kappa}+\frac{d}{\mathrm{2}}}}}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathrm{for}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathit{\kappa}>\mathrm{0}\\ \text{(2b)}& {\displaystyle}& {\displaystyle}\propto {m}^{-\mathrm{2}\mathit{\kappa}-d}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathrm{for}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}m\gg {a}^{-\mathrm{1}}.\end{array}$$

The PSDF obeys a power-law decay at large wavenumbers: $P\left(m\right)\propto {m}^{-\mathrm{2}\mathit{\kappa}-\mathrm{3}}$
for the 3-D case and $P\left(m\right)\propto {m}^{-\mathrm{2}\mathit{\kappa}-\mathrm{1}}$ for the 1-D case, where *κ*
corresponds to the Hurst number. In the following, we will basically use a von
Kármán-type function for characterizing the earth-medium heterogeneity.

Especially for an anisotropic case, we define the von Kármán-type PSDF in 3-D (e.g., Wu et al., 1994; Nakata and Beroza, 2015):

$$\begin{array}{}\text{(3)}& {\displaystyle}P\left(\mathit{m}\right)={\displaystyle \frac{{\mathrm{2}}^{\mathrm{3}}{\mathit{\pi}}^{\frac{\mathrm{3}}{\mathrm{2}}}\mathrm{\Gamma}(\mathit{\kappa}+\frac{\mathrm{3}}{\mathrm{2}}){\mathit{\epsilon}}^{\mathrm{2}}\phantom{\rule{0.125em}{0ex}}{a}_{x}{a}_{y}{a}_{z}}{\mathrm{\Gamma}\left(\mathit{\kappa}\right){\left(\mathrm{1}+{a}_{x}^{\mathrm{2}}{m}_{x}^{\mathrm{2}}+{a}_{y}^{\mathrm{2}}{m}_{y}^{\mathrm{2}}+{a}_{z}^{\mathrm{2}}{m}_{z}^{\mathrm{2}}\right)}^{\mathit{\kappa}+\frac{\mathrm{3}}{\mathrm{2}}}}}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathrm{for}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathit{\kappa}>\mathrm{0}.\end{array}$$

For a case formally
corresponding to *κ*=0 of the von Kármán-type PSDF, we define the
Henyey–Greenstein type ACF and PSDF in 3-D (Henyey and Greenstein, 1941):

$$\begin{array}{}\text{(4a)}& {\displaystyle}& {\displaystyle}R\left(r\right)={\mathit{\epsilon}}^{\mathrm{2}}{K}_{\mathrm{0}}\left({\displaystyle \frac{r}{a}}\right),\text{(4b)}& {\displaystyle}& {\displaystyle}P\left(m\right)={\displaystyle \frac{\mathrm{2}{\mathit{\pi}}^{\mathrm{2}}{\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{3}}}{(\mathrm{1}+{a}^{\mathrm{2}}{m}^{\mathrm{2}}{)}^{\mathrm{3}/\mathrm{2}}}}\approx \mathrm{2}{\mathit{\pi}}^{\mathrm{2}}{\mathit{\epsilon}}^{\mathrm{2}}{m}^{-\mathrm{3}}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}\mathrm{for}\phantom{\rule{0.125em}{0ex}}\phantom{\rule{0.125em}{0ex}}m\gg {a}^{-\mathrm{1}}.\end{array}$$

Note that parameter *ε*^{2} characterizes *P* but *ε*^{2}≠*R*(0) since
*R*(*r*) diverges as *r*→0.

Gaussian-type ACF and PSDF are also used because they are mathematically tractable.

$$\begin{array}{}\text{(5a)}& {\displaystyle}& {\displaystyle}R\left(r\right)={\mathit{\epsilon}}^{\mathrm{2}}{e}^{-\frac{{r}^{\mathrm{2}}}{{a}^{\mathrm{2}}}},\text{(5b)}& {\displaystyle}& {\displaystyle}P\left(m\right)=\sqrt{{\mathit{\pi}}^{\mathrm{3}}}{\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{3}}{e}^{-\frac{{m}^{\mathrm{2}}{a}^{\mathrm{2}}}{\mathrm{4}}}.\end{array}$$

We plot those PSDFs against wavenumber and ACFs against lag distance in Fig. 1.

3 Measurements of random heterogeneities

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There are several kinds of measurements for estimating statistical parameters
characterizing random media. Here we principally collect measurements supposing a von
Kármán-type function for isotropic randomness; however, we include a few measurements
supposing anisotropic randomness and a Gaussian-type function. On a small scale, the photo-scan method is
applied to small rock samples. Acoustic well logs are available in deep wells drilled in
the shallow crust. When the precise velocity tomography result is available, we can
directly calculate the PSDF. In seismology, the most conventional method is to analyze
seismograms of natural earthquakes or artificial explosions after traveling through the
earth heterogeneity. It is better to focus on MS amplitude envelopes (intensity time
traces) since phases are complex and caused by random heterogeneities. Comparing observed
seismogram envelopes with envelopes synthesized in random media, we can evaluate von
Kármán parameters. For the synthesis, we can use the FD simulations, the RTT with the
Born approximation, and the RTT with the phase screen approximation that is equivalent to
the Markov approximation. For each reported measurement, we enumerate the target region,
data and the method, the measured PSDF as a function of wavenumber *m*, von Kármán
parameters (*κ*, *ε*, *a*), the frequency range, the wavenumber range, and
the reference in Tables 1–3. Note that measurements of
heterogeneity listed in the Tables are by no means the only ones. Especially in
seismological measurements, we estimate the wave number range from the frequency range by
using the typical velocity of the target medium. In the Tables, the parameter values in
parentheses (…) are a priori fixed in the measurement. Then, we plot obtained
PSDFs against wavenumber in Figs. 2–5. When the estimated
parameter value is given by a range, a value in square brackets […] is used to
represent for plotting PSDFs in the figures. Measurement with a label with an asterisk
^{*} is insufficient for plotting the PSDF in the figures.

The photo-scan method uses a scanner to take a picture of the polished flat surface of a
small rock sample
(e.g., Sivaji et al., 2002; Spetzler et al., 2002; Fukushima et al., 2003). For the case of
a granite sample, those papers classified
color images on a straight line into three types of mineral grains: quartz, plagioclase,
and biotite. Assigning a typical velocity *V*_{P} or *V*_{S} to each
mineral grain, they constructed a velocity profile along the line. Then, they estimated
the 1-D PSDF of the velocity fractional fluctuation. They measured 1-D PSDFs of granite
and gabbro samples fixing *κ*=0.5 as R1–R5. Figure 2a shows estimated 1-D PSDFs,
where the wavenumber range is on the order of 1 mm^{−1}. We note that raw 1-D PSDFs in
Figs. 4 and 5 of Fukushima et al. (2003) decay a little slower than those of R4 and
R5 in Fig. 2a, especially at large wavenumbers.

In the case of isotropic randomness, we evaluate the 1-D PSDF from the 3-D ACF along the
*z* axis at $x=y=\mathrm{0}$ as follows:

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}{P}_{\mathrm{1}-D}\left({m}_{z}\right)\equiv \underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\int}}{R}_{\text{3-D}}\left(\mathrm{0},\mathrm{0},z\right){e}^{-i{m}_{z}z}\mathrm{d}z\\ {\displaystyle}& {\displaystyle}=\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\int}}\left[{\displaystyle \frac{\mathrm{1}}{{\left(\mathrm{2}\mathit{\pi}\right)}^{\mathrm{3}}}}\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iiint}}{P}_{\text{3-D}}\left({m}_{x}^{\prime},{m}_{y}^{\prime},{m}_{z}^{\prime}\right){e}^{i{m}_{z}^{\prime}z}\mathrm{d}{\mathit{m}}^{\prime}\right]{e}^{-i{m}_{z}z}\mathrm{d}z\\ \text{(6a)}& {\displaystyle}& {\displaystyle}={\displaystyle \frac{\mathrm{1}}{{\left(\mathrm{2}\mathit{\pi}\right)}^{\mathrm{2}}}}\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iint}}{P}_{\text{3-D}}\left({m}_{x}^{\prime},{m}_{y}^{\prime},{m}_{z}\right)\mathrm{d}{m}_{x}^{\prime}\mathrm{d}{m}_{y}^{\prime}.\end{array}$$

Substituting Eq. (3a) into the above equation, we have

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}{P}_{\mathrm{1}-D}\left({m}_{z}\right)\\ {\displaystyle}& {\displaystyle}={\displaystyle \frac{\mathrm{1}}{{\left(\mathrm{2}\mathit{\pi}\right)}^{\mathrm{2}}}}\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iint}}{\displaystyle \frac{\mathrm{8}{\mathit{\pi}}^{\mathrm{3}/\mathrm{2}}{\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{3}}\mathrm{\Gamma}\left(\mathit{\kappa}+\mathrm{3}/\mathrm{2}\right)}{\mathrm{\Gamma}\left(\mathit{\kappa}\right){\left[\mathrm{1}+{a}^{\mathrm{2}}\left({{m}_{x}^{\prime}}^{\mathrm{2}}+{{m}_{y}^{\prime}}^{\mathrm{2}}+{m}_{z}^{\mathrm{2}}\right)\right]}^{\mathit{\kappa}+\mathrm{3}/\mathrm{2}}}}\mathrm{d}{m}_{x}^{\prime}\mathrm{d}{m}_{y}^{\prime}\\ \text{(6b)}& {\displaystyle}& {\displaystyle}={\displaystyle \frac{\mathrm{2}{\mathit{\pi}}^{\mathrm{1}/\mathrm{2}}\mathrm{\Gamma}\left(\mathit{\kappa}+\mathrm{1}/\mathrm{2}\right){\mathit{\epsilon}}^{\mathrm{2}}a}{\mathrm{\Gamma}\left(\mathit{\kappa}\right){\left(\mathrm{1}+{a}^{\mathrm{2}}{m}_{z}^{\mathrm{2}}\right)}^{\mathit{\kappa}+\mathrm{1}/\mathrm{2}}}}.\end{array}$$

Thus, we can evaluate the 3-D PSDF from the 1-D PSDF using parameters *ε*,
*κ*, and *a* of 1-D PSDF.

Supposing the randomness is isotropic, we evaluate corresponding 3-D PSDFs of R1–R5 and plot them in Fig. 2b.

An acoustic well log is obtained from the measurement of the travel time of an ultrasonic
pulse along the wall of a borehole. Measurements W1 (volcanic tuff) and W2 (tertiary to
pre-tertiary) in Japan clearly show power-law decay with *κ*=0.225 and 0.045,
respectively; however, a corner is not clearly seen in each PSDF. Measurement W4 at the
deep well KTB in Germany shows *κ*=0.10. Measurement W3 in the same well shows that
the exponent of wavenumber is −0.97, which formally corresponds to a negative *κ*.
Measurement W5 at Cajon Pass in California shows *κ*=0.11. All these measurements
show very small *κ* values close to 0. Shiomi et al. (1997) made a list of
reported exponents of wavenumber, which shows that most *κ* values are smaller than
0.25. Measurement of *a* seems to be restricted by the sample length. We enumerate those
measurements in Table 1 and plot their 1-D PSDFs against wavenumber in
Fig. 2a. Figure 2b plots the corresponding 3-D PSDFs of W4 and W5.

We note that Wu et al. (1994) measured anisotropy of randomness from the analysis of well logs obtained from two parallel wells at KTB: the ratio of characteristic scales in horizontal to vertical directions ${a}_{h}/{a}_{z}=\mathrm{1.8}$ (see Eq. 3) as shown in W3.

Nakata and Beroza (2015)da Silva et al. (2018)Rothert and Ritter (2000)Gaebler et al. (2015)Przybilla et al. (2009)Calvet and Margerin (2013)Sens-Schönfelder et al. (2009)Sens-Schönfelder et al. (2009)Emoto et al. (2017)Takemura et al. (2009)Kobayashi et al. (2015)Takahashi et al. (2009)Saito et al. (2005)Sanborn et al. (2017)Sanborn et al. (2017)Wang and Shearer (2017)Wang and Shearer (2017)Petukhin and Gusev (2003)Hock et al. (2004)Hock et al. (2004)Yoshimoto et al. (1997b)Takahashi et al. (2009)Morioka et al. (2017)Sens-Schönfelder et al. (2009)Furumura and Kennett (2005)Aki (1973)Capon (1974)Powell and Meltzer (1984)Flatté and Wu (1988)There have been measurements of velocity tomography at various scales, from which we can
calculate the PSDF and then estimate von Kármán-type parameters. This method depends
on the spatial resolution of the tomography result. Measurement L1 in Table 2
is calculated from the precise *V*_{P} tomography result of the shallow crust, Los
Angeles, California: the exponent of wavenumber is −3.08 (*κ*=0.04). Anisotropic
randomness is also reported: *a*_{z}=0.1 and *a*_{h}=0.5 km (see Eq. 3), we
show those in Fig. 3a. Measurement M2 in Table 3 is evaluated
from the 2-D PSDF of the *V*_{S} tomography result of the upper mantle in a low
wavenumber range. Although there is a resolution limit of the tomography method, the
exponent of wavenumber is between −2 and −3, which means $\mathrm{0}<\mathit{\kappa}<\mathrm{0.5}$. We note that
Fig. 8 of Mancinelli et al. (2016a) shows that the 1-D PSDF estimated from the
*V*_{P} tomography result in the upper mantle (Meschede and Romanowicz, 2015)
covers that of MU2 (*κ*=0.05, *ε*=0.1, *a*=2000 km) for the wavenumber
range from $\mathrm{2}\times {\mathrm{10}}^{-\mathrm{4}}$ to 10^{−2} km^{−1}.

Teleseismic *P* waves registered by a large aperture array were used for the evaluation
of the 3-D PSDF of the lithosphere beneath the array: LA1 and LA2 of Table 2 in
Montana and LA3 in southern California used amplitude and phase coherence analyses, where
a Gaussian-type PSDF (Eq. 5b) was assumed because of mathematical simplicity.
As shown in Fig. 3b, they drop very fast as wavenumber increases. Later
Flatté and Wu (1988) developed the angular coherence analysis in addition to the above
methods. Analyzing teleseismic *P* waves registered at NORSAR, they proposed an overlapping two-layer
model LA4, which is composed of a band-limited flat spectrum from the surface to 200 km
of depth and *m*^{−4} spectrum (*κ*=0.5, *ε*=0.01–0.04) for depths from 15
to 250 km. It means *κ*<0.5 and the decay of their PSDF is much smaller
than that of Gaussian types (not shown in Fig. 3b).

FD simulation is often used for the numerical simulation of waves in an inhomogeneous
velocity structure. For the evaluation of average MS amplitude envelopes, we have to
repeat simulations of the wave propagation through random media having the same PSDFs
that are generated by using different random seeds. There are several measurements of
statistical parameters using FD such as L9–L11 and ML4 in Tables 2 and
3. Measurement LS5 focused on the fact that the subducting oceanic plate is an
efficient waveguide for high-frequency seismic waves: estimated anisotropic parameters
are *κ*=0.5 and *ε*=0.02 with *a*_{p}=10 km and
*a*_{t}=0.5 km in the parallel and transverse directions, respectively. Note that
ML2 supposes a Gaussian-type function.

The RTT is essentially stochastic to directly synthesize the intensity (the average MS amplitude envelope) of a wavelet propagating through random media. There are two conventional methods on the basis of the RTT: one uses the Born approximation and the other uses the phase screen approximation based on the parabolic approximation when the wavenumber is larger than the corner. The former neglects the phase shift, but the latter correctly considers the phase shift.

We here study the deterministic scattering of scalar waves by a single spherical obstacle
(radius *a*=5 km and velocity anomaly $\mathit{\epsilon}=+\mathrm{0.05}$) embedded in a homogeneous
medium (*V*_{0}=4 km s^{−1}) as a mathematical model. The Born approximation calculates
spherically outgoing scattered waves putting the incident plane wave of wavenumber
*k*_{c} in the interaction term of the first-order perturbation equation. From the
scattering amplitude we evaluate the total scattering cross section *σ*_{0} as a
measure of scattering power of the obstacle. The resultant *σ*_{0} monotonously
increases with frequency as shown by a blue line in Fig. 6. As
the wavenumber increases (*a**k*_{c}≫1), the phase shift increases as the
incidence plane wave penetrates into
the obstacle. Putting the phase modulated wave according
to the parabolic approximation (the phase screen approximation) into the interaction term
of the first-order perturbation equation, we calculate the scattering amplitude and then
the total scattering cross section. This is the distorted-wave Born approximation with
the phase screen approximation, which is also referred to as the eikonal
approximation.
This approximation predicts that *σ*_{0} (a red line in
Fig. 6) saturates at high frequencies and converges to 2*π**a*^{2}, which is twice the geometrical section area of the obstacle as predicted by shadow
scattering (e.g., Landau and Lifshitz, 2003, p. 519 and 543). We recognize that the
conventional Born approximation is still accurate even for *a**k*_{c}>1; however, it works well only for ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}\mathit{\lesssim}O\left(\mathrm{0.1}\right)$. We should use the distorted-wave Born approximation with the phase
screen approximation for ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}\mathit{\gtrsim}O\left(\mathrm{1}\right)$. The
two approximations predict nearly the same *σ*_{0} value in the
intermediate range. We note that 2*ε**a**k*_{c} is the phase
shift on the center line after passing the obstacle. Note that the phase
screen approximation is not applicable for *a**k*_{c}<1 since it is
based on the parabolic approximation.

Interpreting *ε* and *a* as the RMS fractional fluctuation and the
characteristic length of uniformly distributed random media, we may use the
inequality ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}\ll O\left(\mathrm{1}\right)$ or ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}\mathit{\lesssim}O\left(\mathrm{0.1}\right)$ as a criterion of the Born approximation used
in the RTT.

For uniformly distributed random media characterized by *P*(*m*), the Born
approximation leads to the scattering coefficient at wavenumber
*k*_{c} into scattering angle *ψ*:

$$\begin{array}{}\text{(7a)}& {\displaystyle}g({k}_{\mathrm{c}},\mathit{\psi})={\displaystyle \frac{{k}_{\mathrm{c}}^{\mathrm{4}}}{\mathit{\pi}}}P\left(\mathrm{2}{k}_{\mathrm{c}}\mathrm{sin}{\displaystyle \frac{\mathit{\psi}}{\mathrm{2}}}\right),\end{array}$$

which is axially symmetric. The total scattering coefficient is

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}{g}_{\mathrm{0}}\left({k}_{\mathrm{c}}\right)\equiv {\displaystyle \frac{\mathrm{1}}{\mathrm{4}\mathit{\pi}}}\oint g({k}_{\mathrm{c}},\mathit{\psi})\mathrm{d}\mathrm{\Omega}={\displaystyle \frac{\mathrm{1}}{\mathrm{2}}}\underset{\mathrm{0}}{\overset{\mathit{\pi}}{\int}}g({k}_{\mathrm{c}},\mathit{\psi})\mathrm{sin}\mathit{\psi}\mathrm{d}\mathit{\psi}\\ \text{(7b)}& {\displaystyle}& {\displaystyle}=\underset{\mathrm{0}}{\overset{\mathrm{2}{k}_{\mathrm{c}}}{\int}}{g}_{\mathrm{ker}}({k}_{\mathrm{c}},m)\mathrm{d}m,\end{array}$$

where $m=\mathrm{2}{k}_{\mathrm{c}}\mathrm{sin}\frac{\mathit{\psi}}{\mathrm{2}}$. The integral kernel in the wavenumber space is given by

$$\begin{array}{}\text{(7c)}& {\displaystyle}{g}_{\mathrm{ker}}({k}_{\mathrm{c}},m){\displaystyle}={\displaystyle \frac{{k}_{\mathrm{c}}^{\mathrm{2}}}{\mathrm{2}\mathit{\pi}}}mP\left(m\right).\end{array}$$

The upper bound of the integral is twice the wavenumber.
As an example, Fig. 7 shows plots of *P*(*m*) (blue) vs. *m*, and
*g*_{ker}(*m*) vs. *m* at 0.1 Hz (red) and 1 Hz (green) for the case of
*κ*=0.5, *ε*=0.05, *a*=1 km, and *V*_{0}=4 km s^{−1}. As
shown at the upper-right corner, the scattering pattern at 1 Hz (green)
has a large lobe into
the forward direction; however, it becomes isotropic as the frequency decreases. Dots on
each *g*_{ker} curve show corresponding scattering angles.

In the framework of the RTT, the Monte Carlo simulation is a simple method to
stochastically synthesize the wavelet intensity time trace. A particle carrying unit
intensity is shot randomly from a point source, and its trajectory is traced with the
increment of time steps. The occurrence of scattering is stochastically tested by the
inequality *g*_{0}*V*_{0}Δ*t*> Random[0,1] at every time step of Δ*t*, and
$g({k}_{\mathrm{c}},\mathit{\psi})/\left(\mathrm{4}\mathit{\pi}{g}_{\mathrm{0}}\right({k}_{\mathrm{c}}\left)\right)$ is used as the probability of
scattering into angle *ψ*. Note that Random[0,1] is a uniform random variable between
0 and 1. Since *g*_{0}*V*_{0}Δ*t* is chosen to be small enough, scattering does not occur
every time step but intermittently. As a simple example, Fig. 8a
schematically illustrates the flowchart of the Monte Carlo simulation for the isotropic
radiation from a point source in uniform random media. At lapse time *t*, dividing the
number of particles *n* registered in a spherical shell of radius *r* and a thickness
Δ*r* by the total number of particles *N* and the shell volume 4*π**r*^{2}Δ*r*, we calculate the intensity Green function *G*(*r*,*t*). The intensity time trace *I*(*r*,*t*)
is calculated by the convolution of *G*(*r*,*t*) and the source intensity time function
*S*(*t*) in the time domain. It is easy to introduce a layered structure of background
velocity and intrinsic absorption into the simulation code.

The RTT for the scalar wave case can be extended to the elastic vector wave case by using
Stokes parameters. There are four scattering modes: PP, PS, SS, and SP scatterings, and
the *S*-wave scattering coefficients are not axially symmetric
(see Sato et al., 2012, Fig. 4.7). Many papers
(e.g., Shearer and Earle, 2004; Przybilla et al., 2009) suppose proportional relations
$\mathit{\delta}{V}_{\mathrm{p}}/{V}_{\mathrm{P}\mathrm{0}}=\mathit{\delta}{V}_{\mathrm{S}}/{V}_{\mathrm{S}\mathrm{0}}=\mathit{\xi}$ and $\mathit{\delta}\mathit{\rho}/{\mathit{\rho}}_{\mathrm{0}}=\mathit{\nu}\phantom{\rule{0.125em}{0ex}}\mathit{\xi}$ based on the empirical Birch's law, which reduce three fractional
fluctuations into one (e.g., Sato et al., 2012, Eqs. 4.58 and 4.59).

The RTT with the Born approximation has been often used not only for the analyses of *S*
coda envelopes but also for the whole seismogram envelope from the *P* onset via *P* coda
through *S* wave until *S* coda (see Tables 2 and 3). This method has
been often used not only for the analyses of regional seismograms propagating through the
lithosphere but also for the analyses of teleseismic waves propagating through the
mantle. This method is not only applied to direct *P* phase waves but also PcP and
PKP_{prec}
phase waves and so on. In this review, we neglect intrinsic attenuation
parameters a priori assumed or measured in each paper. For a given wavenumber range
(*k*_{l},*k*_{u}) (gray) in Fig. 7, each PSDF curve using this method in
Figs. 3–5 and 9 is drawn by a dotted line for
(0,*k*_{l}) and a solid line for (*k*_{l},2*k*_{u}) as the line next to the bottom of
Fig. 7. As indicated by dots on the *g*_{ker} curves, the
wavenumber interval of the solid line reflects wide-angle scattering and that of dotted
line reflects narrow-angle scattering around the forward direction.

Most measurements of *S* waves in the lithosphere cover the wavenumber range up to
100 km^{−1}. Measurement L2 analyzed cross-hole seismograms on the order of kHz by
using 2-D RTT, of which the wavenumber range
is as high as on the order of 1 m^{−1}. Measurement MU2 for teleseismic *P*-wave
envelopes at long periods in the upper mantle shows that the characteristic scale
*a*=2000 km is much larger than those of MU3 and ML1 at short periods. Several
measurements a priori suppose *κ*=0.5; however, most measurements show *κ*<0.5
except L6. Measurements ML3 and MW1 propose the HG type (see Eq. 4b)
corresponding to *κ*=0 for the lower or whole mantle. We note that
Mancinelli et al. (2016b) proposed an alternative model of 3-D PSDF
$\propto {m}^{-\mathrm{2.6}}$ in addition to
ML3 (not shown in Fig. 4).

When *a**k*_{c}≫1, scattering mostly occurs within a narrow angle around the
forward direction. At a large travel distance, the wavelet just after the onset is mostly
composed of those waves. The phase screen approximation correctly calculates the phase
shift modulation. For the incidence of a plane wave into the *z* direction, the mutual
coherence function (MCF) of the phase shift modulated waves for an increment Δ*z*
is given by

$$\begin{array}{}\text{(8a)}& {\displaystyle}{\displaystyle}\mathrm{\Phi}({k}_{\mathrm{c}},{r}_{\perp},\mathrm{\Delta}z)={e}^{-{k}_{\mathrm{c}}^{\mathrm{2}}\left(A\right(\mathrm{0})-A({r}_{\perp}\left)\right)\mathrm{\Delta}z}.\end{array}$$

The longitudinal integral of the ACF is

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}A\left({r}_{\perp}\right)=\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\int}}R({\mathit{x}}_{\perp},z)\mathrm{d}z\\ \text{(8b)}& {\displaystyle}& {\displaystyle}={\displaystyle \frac{\mathrm{1}}{(\mathrm{2}\mathit{\pi}{)}^{\mathrm{2}}}}\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iint}}P({\mathit{m}}_{\perp},{m}_{z}=\mathrm{0}){e}^{i{\mathit{m}}_{\perp}{x}_{\perp}}\mathrm{d}{\mathit{m}}_{\perp},\end{array}$$

where *x*_{⊥} is the transverse coordinate vector and ${r}_{\perp}=\left|{\mathit{x}}_{\perp}\right|$ Sato et al. (2012, Eq. 9.60). Taking the Fourier
transform of MCF Φ with respect to transverse coordinates, we have

$$\begin{array}{ll}{\displaystyle}& {\displaystyle}\stackrel{\mathrm{\u02d8}}{\mathrm{\Phi}}({k}_{\mathrm{c}},{k}_{\perp},\mathrm{\Delta}z)={\displaystyle \frac{\mathrm{1}}{(\mathrm{2}\mathit{\pi}{)}^{\mathrm{2}}}}\underset{-\mathrm{\infty}}{\overset{\mathrm{\infty}}{\iint}}\mathrm{\Phi}({k}_{\mathrm{c}},{r}_{\perp},\mathrm{\Delta}z){e}^{i{k}_{\perp}{x}_{\perp}}\mathrm{d}{\mathit{x}}_{\perp}\\ \text{(8c)}& {\displaystyle}& {\displaystyle}\underset{\mathrm{\Delta}z\to \mathrm{0}}{\u27f6}\mathit{\delta}\left({\mathit{k}}_{\perp}\right).\end{array}$$

Since ${\iint}_{-\mathrm{\infty}}^{\mathrm{\infty}}\stackrel{\mathrm{\u02d8}}{\mathrm{\Phi}}({k}_{\mathrm{c}},{k}_{\perp},\mathrm{\Delta}z)\mathrm{d}{k}_{\perp}=\mathrm{1}$, interpreting $\stackrel{\mathrm{\u02d8}}{\mathrm{\Phi}}({k}_{\mathrm{c}},{k}_{\perp},\mathrm{\Delta}z)$ as the probability of ray-bending angle $\mathit{\psi}={\mathrm{tan}}^{-\mathrm{1}}\frac{{k}_{\perp}}{{k}_{\mathrm{c}}}$
per increment Δ*z*=*V*_{0}Δ*t*, we can stochastically synthesize the intensity by
using the Monte Carlo simulation (e.g., Williamson, 1972; Takahashi et al., 2008; Saito et al., 2008). As a simple example, Fig. 8b schematically
illustrates the flowchart of the RTT with the phase screen approximation for the
isotropic radiation from a point source in uniform random media. Different from the Born
approximation, narrow-angle ray bending occurs at every time step. The intensity Green
function can be obtained in the same manner as the RTT with the Born approximation. This
approximation synthesizes the intensity time trace having a delayed peak from the onset
and a decaying tail of early coda at large travel distances. This approximation can not
synthesize the late coda intensity since wide-angle scattering is neglected. The Markov
approximation is known as a stochastic extension of the phase screen method for the
two-frequency mutual coherence function (e.g., Saito et al., 2002). If we focus on
the intensity time trace around the peak arrival, the Markov approximation and the RTT
with the phase screen approximation show good coincidence
(see Sato and Emoto, 2018, Fig. 8).

When this approximation is used, ${k}_{\mathrm{c}}\gg {a}^{-\mathrm{1}}$ is a priori supposed. Most
of this type of measurement reads the peak delay and the envelope width of each
seismogram envelope. There is some merit to the fact that the peak delay measurement is
rather insensitive to intrinsic absorption. In NE Japan, a *κ* value beneath a
volcano LS2 is smaller than those in the fore-arc side L12 and L13. Note that
narrow-angle scattering around the forward direction dominates in teleseismic wavelets
even if the Born approximation is used for the analysis. Narrow-angle scattering is
mostly produced by the PSDF in low wavenumbers compared with *k*_{c}. For a given
wavenumber range (*k*_{l},*k*_{u}) (gray) in Fig. 7, each PSDF curve using
this method in Fig. 3 is drawn by a dotted line for (0,*k*_{l}) and a solid
line for (*k*_{l},*k*_{u}) as the bottom line of Fig. 7.

Some measurements a priori assumed *κ*=0.5; however, most of measurements report
*κ*<0.5. In the mantle, *κ* is very small and close to zero, and a HG-type
function is also proposed. The RMS fractional fluctuation *ε* is on the order
of 0.1 for rock samples and well-log data, and in the range from 0.01 to 0.1 in the
lithosphere and the upper mantle. Large values are reported for the shallow crust L16 and
beneath a volcano LS3; however, smaller values are reported for the lower mantle. The
characteristic scale *a* varies a lot depending on measurements. The corner wavenumber
*a*^{−1} is not clearly seen in PSDFs of acoustic well logs. Some measurements report
anisotropy: W3 of well-logs, L1 of velocity tomography in the shallow crust, and LS5 in
the subducting oceanic slab. The characteristic length in the vertical direction is
smaller than the horizontal direction in the shallow crust, and that in the transverse
direction is smaller than that in the direction parallel to the subducting slab.

Plotting PSDFs against wavenumber is more informative for understanding the random
heterogeneity compared with enumerating statistical parameter values.
Figure 5 shows the plot of 3-D PSDF vs. wavenumber for all the data covering
a wide wavenumber range (10^{−4}–10^{8} km^{−1}). We recognize that the
Gaussian type PSDFs show a very different behavior from others, which suggests the
Gaussian-type assumption is inappropriate. PSDFs in the lower mantle take smaller values,
and those for volcanoes and for the shallow crust take larger values than others.

Eliminating data supposing a Gaussian-type data LA1–LA3, strong heterogeneity data
LS1–LS4, anisotropy-type data L1 and LS5, and the lower mantle and the whole mantle data
ML1–ML4 and MW1–MW2 from Fig. 5, we plot the rest of the data for the crust
and the upper mantle in Fig. 9. Most *ε* values are in the range
of 0.01–0.1, most *κ* values are less than or equal to 0.5 and many of them are
close to 0, and the high wavenumber end of the power-law decay branch of each PSDF is not
far from each corner wavenumber.

We draw a power-law decay line PSDF$\left(m\right)=\mathrm{0.01}{m}^{-\mathrm{3}}$ km^{3} (gray) visually fitting to
most PSDF envelopes for a very wide range of wavenumbers (10^{−3}–10^{8} km^{−1}).
This line is not the average of PSDFs. This line looks like an extension of MU2 in the
upper mantle into higher wavenumbers of the shallow crust.

4 Discussions

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It will be necessary for us to measure the small-scale randomness of
sedimentary rock samples. More measurements are necessary in the wavenumber
range 10^{3}–10^{5} km^{−1} since there are few measurements.

In most PSDF measurements, each power-law decay branch is short since the Born
approximation senses the spectrum up to twice the wavenumber. It will be necessary to
measure how each power-law decay branch varies with wavenumber increasing. It will be
necessary to estimate the corner *a*^{−1} in each measurement with a wide wavenumber
range covering, sufficiently large enough, both sides of the corner. The flat part, the
low-wavenumber side of each PSDF drawn by a dotted line in figures is also important as
the cause of narrow-angle scattering.

Although most measurements used in this review analyzed intrinsic attenuation, we did not enumerate them in this review since different assumptions were used in different measurements. It will be necessary for us to systematically measure the PSDF of random heterogeneity in conjunction with intrinsic attenuation.

We should note that there are large variations in *δ*ln *V*_{S}∕*δ*ln *V*_{P} and $\mathit{\nu}\equiv \mathit{\delta}\mathrm{ln}\mathit{\rho}/\mathit{\delta}\mathrm{ln}{V}_{\mathrm{S}}$ in the earth.
Koper et al. (1999) estimated *δ*ln *V*_{S}∕*δ*ln *V*_{P} to be
in the range 1.1–1.5 in the Tonga Slab. Romanowicz (2001) estimated *δ*ln *V*_{S}∕*δ*ln *V*_{P} to be larger than 2.5 in the lower mantle at larger
length scales. Many measurements reported use *ν*=0.8 for the synthesis, which is
appropriate for the shallow lithosphere. Parameter *ν* takes smaller values of 0.17
for volcanic-tuff (Shiomi et al., 1997) and 0.31–0.33 for sandstone and shale
(Kenter et al., 2007). In the mantle, Karato (2008) estimated *ν*=0.23–0.42
for the *S*-wave velocity predicted from the temperature derivatives of seismic wave
velocities and thermal expansion, and *ν*=0.15–0.23 including the influence of
anelasticity. It will be necessary to introduce realistic *δ*ln *V*_{S}∕*δ*ln *V*_{P} and *δ*ln *ρ*∕*δ*ln *V*_{S} in the synthesis.

Figure 9 summarizes reported PSDF measurements supposing isotropic randomness; however, there are measurements clearly showing anisotropic randomness such as W3 and L1 for the shallow crust and LS5 for the oceanic slab. Those may reflect the effect of gravity for the creation of anisotropy. It will be necessary for us to study how a wavelet propagates through anisotropic random heterogeneity of the earth medium (e.g., Margerin, 2006).

In Sect. 3.6.1, we mentioned that the conventional Born approximation is
inapplicable and the phase screen approximation is useful when the phase shift becomes
large as the wavenumber increases. In order to avoid the difficulty, taking the center
wavenumber of a wavelet as a reference, Sato and Emoto (2018) proposed to divided the
PSDF into two components (see also Sato, 2016; Sato and Emoto, 2017). They use the Born
and phase-screen approximations to the short-scale (high wavenumber) and long-scale (low
wavenumber) components, *P*_{S} and *P*_{L}, respectively, in the RTT in
order to simultaneously explain the envelope broadening just after the onset and the
excitation of late coda waves. Figure 10 illustrates the flowchart of
their Monte Carlo simulation. Their spectrum division method looks like an implementation
of the distorted-wave Born approximation in the RTT since it describes wide-angle
scattering for the incidence of the phase-shift modulated wave. They successfully
synthesized intensity time traces that explain FD simulation results for the case of
*a**k*_{c}=23.6 and ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}=\mathrm{1.39}$. It would be
interesting to see how this method may be extended to polarized elastic waves.

We note that some papers numerically show that the RTT with the Born approximation works
well in some cases over the above limitation. Przybilla et al. (2006) synthesizes
vector-wave intensity that fits to that of the FD simulation in 2-D, even for *S* waves
of *a**k*_{c}=58 and ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}=\mathrm{8.4}$ (see their Table 1) if
the wandering effect is convolved as a result of the travel time fluctuation.
Emoto and Sato (2018) show that the synthesized scalar intensity by the RTT with the
Born approximation fits to that of the FD simulation in 3-D, even for the case of
*a**k*_{c}=23.6 and ${\mathit{\epsilon}}^{\mathrm{2}}{a}^{\mathrm{2}}{k}_{\mathrm{c}}^{\mathrm{2}}=\mathrm{1.39}$ when the wandering
effect is convolved. If the earth heterogeneity is represented by a power-law decay power
spectrum for such a wide wavenumber range, which means that the corner wavenumber is very
low, we should carefully examine the applicability of the Born approximation in the RTT.

Acoustic well-log and photo-scan methods faithfully measure the inhomogeneous elastic coefficients. The RTT used here supposes the scattering contribution of random inhomogeneity of elastic coefficients only; however, observed seismograms do not only reflect those but also the scattering contribution of pores and cracks distributed over the earth medium. It will be necessary for us to study their contribution in the intensity synthesis.

In observation, we may take the power-law spectral envelope as a reference curve for
studying the regional differences, especially in the power-law decay part of the PSDF.
The characteristic length *a* seems to increase as the wavenumber of a wavelet decreases
or as the dimension of measurement system becomes large. It reminds us that the
characteristic scale of the slip distribution increases with increasing source dimension
as Mai and Beroza (2002) analyzed finite-fault source inversion results (see their
Fig. 12).

The power-law decay spectral envelope reminds us of the observed fractal nature of
various kinds of surface topographies: Sayles and Thomas (1978a, b)
show a
1-D PSDF $\propto {m}^{-\mathrm{2}}$ for wavelengths 10^{−6}–10^{3} km although the power
exponent varies from −1.07 to −3.03 for small segments; Brown and Scholz (1985)
show a
1-D PSDF $\propto {m}^{-\mathrm{1.64}\phantom{\rule{0.125em}{0ex}}\mathrm{to}\phantom{\rule{0.125em}{0ex}}-\mathrm{3.36}}$ for the wavenumber range
10^{−6} to 0.1 µm^{−1}, especially for the topography of natural rock
surfaces and faults.
We also note that the PSDF of the refractive index fluctuation of air obeys the Kolmogorov spectrum:
3-D PSDF $\propto {m}^{-\mathrm{11}/\mathrm{3}}$, where $\mathit{\kappa}=\mathrm{1}/\mathrm{3}$.
This spectrum is physically produced by the cascade in the turbulent flow of low
viscosity air: the large eddies breaks up originating smaller eddies dissipating energy
by viscosity. However, it may be difficult to apply this cascade model to the mantle
since the viscosity of mantle fluid is thought to be high.

For igneous rocks such as granite, there are variations in composition of minerals and grain sizes, which depend on a variety of slow crystallization differentiations of basaltic magma. Random variations in acoustic well-log profiles reflect the complex sedimentation process during the geological history. Volcanism produces more heterogeneous structures composed of pyroclastic material and lava. For random heterogeneities in the mantle, we imagine complex mantle flow. Mancinelli et al. (2016a) referred to a marble cake mantle model (Allègre and Turcotte, 1986) containing heterogeneity mostly composed of basalt and harzburgite at many scales in the upper mantle in order to explain the power-law spectrum. Stixrude and Lithgow-Bertelloni (2007) proposed the velocity variation due to chemical and phase stability at different depths, which is a possible candidate especially for heterogeneity in the vertical direction. If we accept the power-law decay spectrum, we will have to study what kinds of geophysical mechanisms created such random medium spectra at different scales and in different geological environments in the solid earth.

In advance of the measurements based on the RTT for anisotropic scattering presented
here, there have been many measurements of the isotropic scattering coefficient
*g*_{iso} in the world on the basis of the RTT with the isotropic scattering
assumption (e.g., Sato et al., 2012; Yoshimoto and Jin, 2008). The isotropic scattering
model is mathematically tractable, and the multiple lapse-time window analysis
(Fehler et al., 1992; Hoshiba, 1993) has often been used for practical analyses. This
method essentially estimates *g*_{iso} from the ratio of late coda excitation to
the radiated energy irrespective of the envelope broadening. Recent measurements show
that *g*_{iso} decreases with depth
(e.g., Rachman and Chung, 2016; Badi et al., 2009). It will be interesting to plot the
frequency dependence of reported *g*_{iso} for a wide range of frequencies and to
study the relation with the obtained power spectral envelope shown in
Fig. 9.

5 Conclusions

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Recent seismological observations focusing on the collapse of an impulsive wavelet
revealed the existence of small-scale random heterogeneities in the earth medium. The RTT
has often been used for the study of the propagation of wavelet intensities, the MS
amplitude envelopes. For the statistical characterization of the PSDF of random velocity
inhomogeneities in a 3-D space, we have used von Kármán type functions with three
parameters: the RMS fractional velocity fluctuation *ε*, the characteristic
length *a*, and the order *κ* of the modified Bessel function of the second kind.
This model leads to the power-law decay of PSDF $\propto {m}^{-\mathrm{2}\mathit{\kappa}-\mathrm{3}}$ at wavenumber
*m* higher than the corner at *a*^{−1}. We have compiled reported statistical parameters
of the lithosphere and the mantle based on various types of measurements for a wide range
of wavenumbers: photo-scan data of rock samples, acoustic well-log data, and envelope
analyses of cross-hole experiment seismograms, regional seismograms, and teleseismic
waves based on the RTT. Reported *κ* values are distributed between 0 and 0.5 (PSDF
$\propto {m}^{-\mathrm{3}\phantom{\rule{0.125em}{0ex}}\mathrm{to}\phantom{\rule{0.125em}{0ex}}-\mathrm{4}}$), where many of them are close to 0 (PSDF $\propto {m}^{-\mathrm{3}}$). Reported *ε* values are on the order of 0.01–0.1 in the crust and
the upper mantle, where smaller values in the lower mantle and higher values beneath
volcanoes. Reported *a* values distribute very widely; however, each one seems to be
restricted by the dimension of the measurement system or the sample length. In order to
grasp the spectral characteristics, eliminating strong heterogeneity data and the lower
mantle data, we have plotted all the reported PSDFs in the crust and the upper mantle
against wavenumber *m* for a wide range (10^{−3}–10^{8} km^{−1}). We find that the
envelope of those PSDFs is well approximated by a power-law decay curve
0.01 *m*^{−3} km^{3}. Multiple plots of PSDFs and the power-law decay spectral
envelope obtained require us to do the following. In theory, it will be necessary to
examine whether the Born approximation is reliable or not if the wavenumber increases
much larger than the corner; in observation, we will have to more carefully examine the
behavior of each PSDF on both sides of the corner. If we accept the power-law decay
spectral envelope, it suggests that the earth-medium randomness has a broad spectrum. We
may consider the obtained power-law decay spectral envelope as a reference for studying
the regional differences. It is interesting to study what kinds of geophysical processes
created the power-law spectral envelope at different scales and in different geological
environments of the solid earth medium.

Data availability

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Data availability.

All the data for this review work are listed in Tables 1–3, where all the references are given.

Author contributions

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Author contributions.

The author HS reviewed reported measurement of statistical parameters characterizing the random heterogeneities in the earth.

Competing interests

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Competing interests.

The author declares that they have no conflict of interest.

Acknowledgements

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Acknowledgements.

This paper is a part of the Beno Gutenberg medal lecture at the 2018 EGU assembly held in
Vienna. The author is grateful to the seismology section of EGU for giving him an
opportunity for reviewing measurements of random heterogeneities in the solid earth and
various theoretical approaches. The author expresses sincere thanks to his colleagues and
ex-graduate students who enthusiastically collaborated with him in studying seismic wave
scattering in random media. The author acknowledges reviewers, Vernon Cormier, Michael
Korn, and Ludovic Margerin and editor Tarje Nissen-Meyer for their helpful comments and
suggestions.

Edited by: Tarje Nissen-Meyer

Reviewed by: Ludovic Margerin, Michael Korn, and Vernon Cormier

References

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Short summary

Recent seismological observations clarified that the velocity structure of the crust and upper mantle is randomly heterogeneous. I compile reported power spectral density functions of random velocity fluctuations based on various types of measurements. Their spectral envelope is approximated by the third power of wavenumber. It is interesting to study what kinds of geophysical processes created such a power-law spectral envelope at different scales and in different geological environments.

Recent seismological observations clarified that the velocity structure of the crust and upper...

Solid Earth

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