submitted to Geophys. J. Int.
Transient gravity perturbations induced by earthquake
J. Harms1,2,7, J.-P. Ampuero3, M. Barsuglia4, E. Chassande-Mottin4,
J.-P. Montagner5, S. N. Somala3and B. F. Whiting4,6
a degli Studi di Urbino “Carlo Bo”, Urbino, 61029, Italy
2INFN Sezione di Firenze, Sesto Fiorentino, 50019, Italy
3Seismological Laboratory, California Institute of Technology, Pasadena, CA 91125, USA
4AstroParticule et Cosmologie (APC), CNRS: UMR7164-IN2P3-Observatoire de Paris-
e Denis Diderot-Paris 7, CEA: DSM/IRFU, France
epartement de Sismologie, Institut de Physique du Globe, 75005 Paris, France
6University of Florida, Gainesville, FL 32611, USA
7After the ﬁrst name, the author list is in alphabetic order.
The static and transient deformations produced by earthquakes cause density perturba-
tions which, in turn, generate immediate, long-range perturbations of the Earth’s grav-
ity ﬁeld. Here, an analytical solution is derived for gravity perturbations produced by a
point double-couple source in homogeneous, inﬁnite, non-self-gravitating elastic media.
The solution features transient gravity perturbations that occur at any distance from the
source between the rupture onset time and the arrival time of seismic P waves, which
are of potential interest for real-time earthquake source studies and early warning. An
analytical solution for such prompt gravity perturbations is presented in compact form.
We show that it approximates adequately the prompt gravity perturbations generated by
2J. Harms et al
strike-slip and dip-slip ﬁnite fault ruptures in a half-space obtained by numerical simula-
tions based on the spectral element method. Based on the analytical solution, we estimate
that the observability of prompt gravity perturbations within 10 s after rupture onset by
current instruments is severely challenged by the background microseism noise but may
be achieved by high-precision gravity strainmeters currently under development. Our an-
alytical results facilitate parametric studies of the expected prompt gravity signals that
could be recorded by gravity strainmeters.
Key words: gravity perturbation – double couple – time domain – seismic potentials.
Earthquakes generate transient and static deformation, including volumetric deformation and displace-
ment of material interfaces, which modify the spatial distribution of material density. This mass re-
distribution induces changes in the Earth’s gravitational ﬁeld. Permanent gravity changes generated
by static deformation induced by the co-seismic and post-seismic slip of large earthquakes have been
observed by superconducting gravimeters and gravity ﬁeld satellite missions on multiple occasions
(Imanishi et al. 2004; Wang et al. 2012a; Fuchs et al. 2013; Cambiotti & Sabadini 2013). Theoreti-
cal models of these static gravity perturbations have been developed (Okubo 1992; Sun et al. 2009)
and compared to observations (Matsuo & Heki 2011; Wang et al. 2012b). The transient effects of
elasto-gravitational coupling on the Earth’s gravest normal modes are well established and routinely
included in normal mode computations for long-period seismology (Dahlen & Tromp 1998). Transient
perturbations of the gravity ﬁeld concurrent with the passage of seismic waves have been examined
as a broad-band source of so-called Newtonian noise for gravitational wave detectors (Beker et al.
2011; Driggers et al. 2012). Here we address, for the ﬁrst time, the problem of modelling the gravity
perturbations generated by earthquakes at short time scales, including those generated by transient
deformation induced by seismic waves and those occurring during the fault rupture.
Our focus here is on a particular short-term elasto-gravitational effect that has received no at-
tention in the literature. Owing to the long-range and virtually instantaneous (speed-of-light) effect
of gravitational forces, density perturbations lead immediately to global perturbations of the Earth’s
gravity ﬁeld. In particular, the volumetric deformation carried by P waves induces remote gravity per-
turbations even at distances beyond the P wave front. Hence, at any distance from an earthquake source
and its dynamically deforming region, gravity perturbations are expected to be induced even between
the onset time of the rupture and the arrival time of seismic waves. We refer to these as prompt gravity
Transient gravity perturbations induced by earthquake rupture 3
perturbations. If practically measurable, these signals would add a new dimension to real-time seis-
mology that may enhance rapid earthquake source detection and parameter estimation capabilities and
may contribute to earthquake and tsunami early warning. Some analogies may be drawn to prompt
electromagnetic perturbations caused, in principle, by earthquakes via the electrokinetic effect in sat-
urated porous media (Gao et al. 2013) or via the motional induction effect of the Earth’s conductive
crust and magnetic ﬁeld (Gao et al. 2014). Here we develop building blocks for the quantitative anal-
ysis of prompt gravity perturbations, as needed to assess their detectability and information content.
An analysis of their potential contribution to earthquake early warning will be reported elsewhere, as
well as results of a search for prompt gravity signals in superconducting gravimeter recordings of a
While normal mode theory in global seismology accounts for gravitational effects (Takeuchi &
Saito 1972; Dahlen & Tromp 1998), the structure of the prompt gravity perturbation ﬁeld is not im-
mediately evident from it. To facilitate systematic, parametric studies of transient gravity signals, we
develop in section 2 an analytical model of transient Newtonian gravity perturbations from point shear
dislocations in unbounded, uniform, non-self-gravitating elastic media. The model accounts for the ef-
fects of density perturbations induced by both static and transient deformation and provides a complete
gravity time series, before and after the passage of seismic waves. The key result is a compact expres-
sion for prompt gravity perturbations. Its relevance in more realistic settings is assessed in section 3
by comparing the analytical results with results from numerical simulations that include ﬁnite fault
and free surface effects. In section 4, after realizing that prompt gravity perturbations are generally
too weak to be detected by conventional gravimeters within 10 s after rupture onset, we derive ex-
pressions for transient gravity gradients and assess the theoretical potential of high-sensitivity gravity
strainmeters to observe prompt gravity signals.
2 GRAVITY PERTURBATIONS FROM POINT-SHEAR DISLOCATIONS IN INFINITE,
HOMOGENEOUS, NON-SELF-GRAVITATING ELASTIC MEDIA
2.1 Model assumptions and density perturbations
We consider a point shear dislocation in an inﬁnite, elastic and homogeneous medium. The defor-
mation is related to the dislocation source by the seismic wave equation and the Newtonian gravity
potential perturbation is related to the deformation-induced density perturbations by Poisson’s equa-
tion. We ignore the effects of self-gravitation, i.e. we assume that gravity perturbations induced by
deformation do not act back on the deformation and their contribution to the seismic wave equation
is ignored. Self-gravitation effects are signiﬁcant only at periods much longer than 100 s (see p. 142
4J. Harms et al
of Dahlen & Tromp (1998)), whereas our primary interest here is in the short time-scales over which
signiﬁcant prompt gravity perturbations may develop (the P wave travel time from the hypocenter to
distances of up to a few 100 km). We adopt a source-based coordinate system with origin at the loca-
tion of the shear dislocation, z-axis parallel to the slip direction and x-axis perpendicular to the fault
plane. Spherical coordinates r, θ, φ are related to the Cartesian coordinates via x=rsin(θ) cos(φ),
y=rsin(θ) sin(φ),z=rcos(θ), with 0< θ < π, and 0< φ < 2π. The point-shear dislocation
source is represented by a double-couple moment tensor point-source, with the following equivalent
f(r, t) = −M0(t)∂δ(r)
where exand ezare the unit vectors pointing along the corresponding coordinate axes. We assume a
ﬁxed focal mechanism but allow for an arbitrary seismic moment time function M0(t).
The displacement ﬁeld induced by seismic waves is denoted by ξ(r, t). The resulting density
perturbation δρ(r, t)is given by the linearized continuity equation (e. g. equation 3.46 of Dahlen &
δρ(r, t) = −ρ0∇·ξ(r, t)(2)
where ρ0is the unperturbed mass density. It depends on the divergence of the displacement ﬁeld,
which describes volumetric deformations of the medium. From the known expression of the seismic
wave ﬁeld generated by a point double-couple source (Aki & Richards 2009):
δρ(r, t) = RP(θ, φ) ∆(r, t),(3)
RP(θ, φ) = cos(φ) sin(2θ)
is the quadrupolar P-wave radiation pattern, er≡r/r the radial unit vector, and
M0(t−r/α) + r2
where αis the P-wave speed. Because S waves produce no volumetric deformation, density perturba-
tions in inﬁnite, uniform, elastic media are carried exclusively by P waves and hence depend on the
moment time function delayed by the P-wave travel time, M0(t−r/α). The density perturbation is
plotted in Figure 1 as a function of distance and time, for φ= 0 and θ=π/4. The plot shows that in
addition to a lasting density change in the vicinity of the source, a propagating density perturbation is
transported by the compressional waves.
Transient gravity perturbations induced by earthquake rupture 5
Figure 1. Density perturbation ﬁeld as a function of distance and time, for φ= 0,θ=π/4and seismic moment
time function M0tanh(t/τ)Θ(t), where Θ(·)denotes the unit step function. Since the positive and negative
values vary over many orders of magnitude, the color scale is based on a log-modulus transformation (John &
2.2 Prompt gravity perturbation
Prompt gravity perturbations are essentially the exterior gravitational ﬁeld induced by the density per-
turbation outside its spatial support, that is, beyond the P-wave front. Their structure can be determined
by exploiting classical results in potential theory (Jekeli 2007) as follows. The density perturbation
ﬁeld inherits the symmetries of a quadrupole from the P-wave radiation pattern, which features four
lobes of alternating signs. A quadrupole density perturbation generates an exterior gravitational poten-
tial, which has also a quadrupole structure and decays as 1/r3
0, where r0is the distance to the centroid
of the quadrupole.
In this section, we use seismic potentials (of the seismic ﬁelds and alternatively of the seismic
sources) to obtain an explicit expression of the gravity perturbation in homogeneous space valid for
all times. In the appendix, a second method is described, which uses the spherical multipole expansion.
While the method based on potentials is not very intuitive and does not give a direct interpretation of
the effective source of gravity perturbations, it is more elegant and should prove useful also to solving
more complicated calculations in the future, for example concerning gravity perturbations in a half
The perturbation of the Newtonian gravity potential, δψ(r0, t), is determined by the density ﬂuc-
6J. Harms et al
tuations according to Poisson’s equation
∇2δψ = 4πGδρ (6)
where Gis the gravitational constant. In a uniform medium, the density perturbation is governed by
the seismic displacement ﬁeld via equation (2). We represent the seismic ﬁeld in terms of its Lam´
ξ(r, t) = ∇φs(r, t) + ∇×ψs(r, t)(7)
The scalar potential φsgives rise to compressional waves, whereas the vector potential ψsgives rise
to shear waves and also obeys the condition ∇·ψs= 0. Inserting equation (7) into equation (2) and
noting that ∇·∇×ψs= 0, one ﬁnds
and, from equation (6), ∇2δψ =−4πGρ0∇2φs. Hence, considering the anticipated fast (1/r3
of the gravity potential away from the density perturbation region, one ﬁnds for an inﬁnite medium
δψ(r0, t) = −4πGρ0φs(r0, t).(9)
This equivalence between a potential giving rise to a seismic compressional ﬁeld and the correspond-
ing gravity potential perturbation is remarkable. Given the known solution for seismic potentials from
a point force in inﬁnite media (Aki & Richards 2009), one can derive the expression for a double-
couple and rescale it according to equation (9) to obtain the following expression of the perturbed
δψ(r0, t) = GRP(θ0, φ0)
Note that this result does not depend explicitly on the reference density of the medium. Before the
arrival of P waves, when t<r0/α, the ﬁrst term in brackets vanishes, the upper limit of integration of
the second term can be substituted by t, and
du uM0(t−u) =
if M0(t)=0for t < 0. Therefore, the early gravity potential assumes a remarkably simple form
independent of the speed αof compressional waves:
δψ(r0, t) = −RP(θ0, φ0)3G
Transient gravity perturbations induced by earthquake rupture 7
As anticipated, the prompt gravity potential inherits the quadrupole distribution of the P wave radiation
pattern and decays as 1/r3
0. Its relation to the seismic moment is remarkably simple: it is proportional
to the second integral of the seismic moment time function.
The derivation, in particular equation (9), has the unintuitive feature that the prompt gravity po-
tential perturbation appears to emerge from the “acausal” component of the P-wave potential. This
component has actually no physical signiﬁcance in the seismic-wave equation, where its contribution
to the near-ﬁeld seismic waveﬁeld is cancelled out by a similar contribution from the S-wave potential.
We can provide an alternative derivation of the prompt gravity potential perturbation based on
source potentials, which allows further insight into the origin of its simple structure. Taking the second
time derivative of equation (9) and considering the wave equation for the P-wave potential,
where Φ(r, t)is the scalar Helmholtz potential of the equivalent body force representation of a dislo-
cation (e.g. equation 4.9 of Aki & Richards (2009)), we ﬁnd
Making use of equation (8), we get
ψ=−4πG(−α2δρ + Φ).(15)
The ﬁrst term in brackets is a propagative contribution carried by P waves. The second term is long-
ranged. Before the arrival of P waves to the detector, only the second term is non zero and hence the
prompt gravity potential perturbation satisﬁes
Combining spatial derivatives of the point-force potential given e.g. in equation (4.17) of Aki &
Richards (2009), it can be shown that the scalar Helmholtz potential of the double-couple source
given in equation (1) is
Φ(r, t) = M0(t)
Porting this into equation (16) and integrating twice leads to equation (12). This derivation, in par-
ticular equation (16), provides a more elegant proof that the prompt gravity potential perturbation is
proportional to the second integral of M0(t). Also, equation (15) is consistent with equation (10) but
shows more clearly the separation between propagative and long-ranged components of the gravity
Finally, we provide an explicit expression of the prompt perturbation of gravity acceleration, δa=
8J. Harms et al
−∇δψ, which takes the form
δa(r0, t) = 6G
Reﬂecting the long-range and instantaneous nature of gravity changes, these expressions show that the
prompt gravity perturbation extends over arbitrary distances and involves the seismic moment time-
function without P-wave time delay. Its proportionality to the second integral of seismic moment has
important implications. Compared to far-ﬁeld seismic ground displacements, which are proportional to
seismic moment rate, the high-frequency content of prompt gravity potential perturbations is damped.
For self-similar ruptures, whose moment function has an onset proportional to t3, the prompt gravity
potential signal is predicted to start as t5. After the rupture ends and before the P wave arrives, it grows
as t2. At the arrival of P waves, t=r0/α and δa∝t2/r4
0, consistent with the distance decay
of gravity perturbations produced by static earthquake deformation (Okubo 1991). In the calculations
leading to equation (12), it is found that the near-, intermediate- and far-ﬁeld terms of the density
perturbation, represented by the three terms in equation (5), have comparable contributions to the
prompt gravity perturbation. In particular, the contribution of the static deformation (the near-ﬁeld
term) is not dominant at any distance.
3 COMPARISON WITH SPECTRAL ELEMENT SIMULATIONS
The theory developed in section 2 does not capture all effects that may be present in nature. Even if
for some purposes the Earth can be sufﬁciently approximated by a homogenous medium, it remains
to be determined under which conditions this theory adequately describes ﬁnite fault sources in a half
First, the theoretical results are derived for a point source. Additional ﬁnite-source effects are
expected close to a ﬁnite-size rupture. We do not consider this a major drawback, since the problem
is linear and ﬁnite-sized sources of arbitrary shape and with complicated rupture propagation can
be represented by superposition of point sources. A suitable approximation to describe ﬁnite-size
ruptures on a single planar fault is proposed here. The prompt gravity perturbation in that scenario is a
superposition of quadrupole gravity ﬁelds, each of the form given by equation (12). The leading term
of the multipole expansion of a superposition of quadrupoles is also a quadrupole, whose quadrupole
moment is simply the sum of the individual ones. The higher order multipole terms (octupole, etc) are
not necessarily zero but they decay faster with distance. Hence, at a distance from the P-wave front, the
prompt gravity perturbation produced by a ﬁnite earthquake source is approximately given by equation
(12) if, as conventionally in seismology, M0(t)is understood as the seismic moment integrated over
the whole ﬁnite rupture surface. To minimize the octupole contribution, r0should be understood as the
Transient gravity perturbations induced by earthquake rupture 9
distance to the instantaneous earthquake centroid (or, more generally, to the instantaneous centroid of
the density perturbation ﬁeld). The quality of this approximation is expected to degrade as the P-wave
Second, the analytical solution was derived for inﬁnite media. If an event occurs at a depth h, then
we can expect deviations from our results as soon as P waves reach the surface at time h/α, which is
typically a few seconds only. Therefore, simulating many tens of seconds of gravity time series using
the analytical results of section 2 would in general lead to inaccurate predictions. However, for early
warning applications of prompt gravity signals the important question is if the inﬂuence of the surface
is signiﬁcant within a few seconds of the rupture onset.
Finite-source and free-surface effects can be assessed by comparing the analytical prediction
with results obtained from a numerical simulation on a half-space. The simulation tool used here
is SPECFEM3D, a seismic wave propagation code based on the spectral element method (Komatitsch
& Vilotte 1998; Komatitsch & Tromp 1999) which can simulate ﬁnite earthquake sources with pre-
scribed slip rate (S. N. Somala, J.-P. Ampuero and N. Lapusta, Finite-fault source inversion using ad-
joint methods in 3D heterogeneous media, submitted manuscript). We implemented in SPECFEM3D
the computation of gravity acceleration perturbations above the half-space surface produced by the
seismic displacement ﬁeld ξ(r, t)as a function of time, based on the following equation (Harms et al.
(2009); equations 3.100-3.101 of Dahlen & Tromp (1998)),
δa(r0, t) = Gρ0ZdV1
|r−r0|3(ξ(r, t)−3(er·ξ(r, t)) ·er).(19)
This equation has the form of a dipole perturbation associated with displaced point masses ρ0dV. It is
valid as long as the displacement is much smaller than the distance |r−r0|. The integral is performed
over a ﬁxed domain but it accounts for the gravity perturbations induced by deformation of the Earth’s
surface. Equation (19) gives the Eulerian gravity acceleration measurable by a seismically isolated
sensor, whereas the Lagrangian quantity measured by a non-isolated sensor involves an additional term
representing advection through the initial gravity gradient (Dahlen & Tromp 1998). This distinction is,
however, irrelevant before seismic waves arrive to the sensor. In this section, we compare the simulated
half-space gravity perturbation with the full-space analytical solution of prompt gravity acceleration
perturbations, equation (18).
The assumed kinematic source model is speciﬁed in appendix B. It represents a circular pulse-
like rupture of ﬁnite duration τwith cosine slip rate function, uniform ﬁnal slip δ, rise time Tand
constant rupture speed vrup. From the rupture model, an explicit expression of the moment function
M0(t)is obtained that is inserted into equation (18) to calculate an analytical prediction of the gravity
perturbation. We present results of three simulations, a very deep strike-slip event, a shallow one and
10 J. Harms et al
Parameter Symbol Value
Rupture speed vrup 3 km/s
P-wave speed α6.4 km/s
Mass density ρ02670 kg/m3
Shear modulus µ27 GPa
Slip δ0.5 m
Total ruptured area A110 km2
Total seismic moment M01.5×1018 Nm
Moment magnitude Mw6.1
Total ruptured area A400 km2
Total seismic moment M05.4×1018 Nm
Moment magnitude Mw6.5
Table 1. Parameter values used for the comparison between the numerical simulation and theory.
a shallow dip-slip event, and compare them with the analytical predictions. The duration of the strike-
slip events is τ= 2 s, while the duration of the dip-slip event is 4 s. The rise time in all cases is
T= 1 s. We simulated about 5 s of time series, a short time scale relevant to anticipated applications
in early warning. It is computationally expensive to increase the simulation duration and at the same
time increase the size of the simulation domain to encompass the complete P wave front and preserve
accuracy at high frequencies. Further parameter values of the simulations are listed in Table 1. For the
strike-slip scenarios, the total ruptured area is A=π(vrupτ)2and the seismic moment M0=µδA. In
the dip-slip scenario the rupture reaches the surface.
Figure 2 shows the comparison between theory and simulation for the deep strike-slip event with
hypocenter at 50 km depth. We compare the gravity acceleration in the horizontal direction normal to
the fault evaluated at a horizontal distance of 1km to the epicenter and to the fault (to avoid a null of
gravity acceleration) and 2 km above ground (to avoid exponentially decaying artifacts of the gravity
calculation related to the ﬁnite grid density). The relative deviation (shown as dashed line) grows with
time but stays smaller than 7% over the entire 5 s. This serves as veriﬁcation of our SPECFEM3D
implementation and to show that ﬁnite source effects are not signiﬁcant at this distance.
Figure 3-left shows the comparison for the shallow strike-slip event with hypocenter depth 7.5 km.
Gravity is evaluated at 50 km horizontal distance to the epicenter and 2 km above ground. The hori-
Transient gravity perturbations induced by earthquake rupture 11
Figure 2. Comparison between analytical predictions gth and numerical simulation gnum of prompt gravity
acceleration perturbations for a deep strike-slip earthquake. The solid curve is plotted against the right y-axis,
while the dashed and dash-dotted curves are plotted against the left y-axis. The hypocenter is 50 km deep so that
surface effects are non-existent during the ﬁrst 5 s. The detector is located at a horizontal distance of 1 km from
the epicenter to avoid a null of gravity acceleration. The result shows that approximating the fault rupture as a
point source leads to relative deviations smaller than 0.07 within the ﬁrst 5s.
zontal distance was chosen larger than the propagation distance of seismic waves during the ﬁrst 5 s.
The plot in the right of Figure 3 shows the comparison for the 7.5 km deep dip-slip event with 45◦dip
angle, and 4 s duration. In these two examples, the relative deviation is smaller than 1.5% and 6%, re-
spectively. This is even smaller than in the previous example, which may seem surprising given that at
t= 5 s seismic waves have already been reﬂected from the surface or have been converted into surface
waves, which was not considered in the theoretical analysis. Especially concerning the dip-slip event,
one might have expected more signiﬁcant deviations since the surface experiences a differential lift
across the fault trace when the rupture reaches the surface at about t= 3.5s. The agreement between
the theoretical and numerical results even in this case may be taken as a hint of a more fundamental
principle that is to be discovered in the full half-space solution.
Strong deviations between theory and numerical simulations were only observed for shallow
events in cases where gravity was evaluated close to the epicenter so that the contribution from passing
Rayleigh waves was dominant. An open question that should be addressed in future studies is whether
12 J. Harms et al
Figure 3. Comparison between analytical predictions gth and numerical simulation gnum of prompt gravity
acceleration perturbations for a shallow strike-slip earthquake and a surface-breaking dip-slip earthquake. Same
representation as in Figure 2. Left: strike-slip event with 2 s duration. Right: dip-slip event with dip angle of 45◦
and duration of 4 s. In both cases, the event hypocenter is 7.5 km deep so that surface effects can be expected
shortly after 1 s, and the detector is located at 50 km distance from the epicenter so that seismic waves do not
reach it within the 5 s of simulation time. For the dip-slip scenario, the rupture reaches the surface at about 3.5 s.
The impact of the surface onto gravity perturbations stays weak within the ﬁrst 5 s, leading to relative deviations
smaller than 0.015 for the strike-slip event, and 0.06 for the dip-slip event.
results still agree well even long after 5 s, at least as long as contributions from passing Rayleigh waves
can be excluded. The effect of heterogeneity of the crust also deserves further scrutiny.
4 TRANSIENT GRAVITY GRADIENTS
Prompt gravity perturbations produced by earthquakes have not been observed with existing gravime-
ters or seismometers. An order-of-magnitude estimate of the amplitude of the prompt gravity accelera-
tion signal can be obtained from equation (18), neglecting the dependence on directions. We assume a
self-similar rupture model, whose seismic moment function starts as M0(t)∼t3. Based on empirical
relations between seismic moment and rupture duration (Houston 2001) we set M0(1 s) = 1017 Nm.
We ﬁnd that
∼O(10 nm/s2)(t/10 s)5
with t < r0/α (before the P-wave arrival). Note that 50 km is the typical distance travelled by a P wave
in 10 s. This estimate is only valid before the half-duration of the rupture; after that it over-estimates
the signal amplitude. At late times, if the rupture duration and distance are large enough, prompt
gravity perturbations above 10 nm/s2are predicted. However, gravity acceleration changes forming
within the ﬁrst few seconds at local distances to the source (50 km or less) lie well below 10 nm/s2.
Transient gravity perturbations induced by earthquake rupture 13
This is also illustrated in the ﬁgures in Section 3. Measurements of such weak gravity transients are
masked by a foreground of seismic noise (Berger et al. 2004; Brown et al. 2014), which is of order
100 nm/s2between 0.1 Hz and 1 Hz. A novel class of instruments, for example seismically isolated
gravity gradiometers (also referred to as gravity strainmeters; (Harms et al. 2013)), is required to
detect these early transients.
The past two decades have seen rapid progress in the development of ultra-sensitive gravity strain-
meters, mainly driven by scientiﬁc communities working on gravitational-wave (GW) detection from
astrophysical sources. Gravity strain is the relative change in physical distance between two test
masses in response to a changing gravitational ﬁeld. This change can be produced by a GW of as-
trophysical origin as well as by gravity perturbations from terrestrial sources. Laser-interferometric
GW detectors such as LIGO (LIGO Scientiﬁc Collaboration 2009), Virgo (Accadia, T. et. al. 2011),
uck et al. 2010), and KAGRA (Aso et al. 2013) measure gravity strain as differential dis-
placement between seismically isolated test masses using high-power, in-vacuum lasers. Strain sen-
sitivities of better than 10−22Hz−1/2have been demonstrated in the frequency range between about
50 Hz and 1000 Hz. The LIGO and Virgo detectors are currently being upgraded to advanced con-
ﬁgurations with design strain sensitivity better than 10−23Hz−1/2between about 30 Hz and 2000 Hz.
Furthermore, future ground-based detectors have been proposed, such as the Einstein Telescope (ET
Science Team 2011) with even more enhanced strain sensitivity, and a detection band extending down
to a few Hz. In parallel to these kilometer-scale detectors, groups are developing gravity strainmeters
targeting signals below 1 Hz, which are better suited to detect gravity perturbations changing over
timescales of a few seconds (Ando et al. 2010; Hohensee et al. 2011; Dickerson et al. 2013; Shoda
et al. 2014).
In the next section, we derive expressions for gravity strain perturbations (expressed as gravity gra-
dients, i. e. the second time derivative of gravity strain) useful to assess the capability of future sensors
to observe prompt gravity signals. Current concepts and sensitivity goals of these ”low-frequency”
instruments are brieﬂy reviewed in section 4.4.
4.2 Gravity-gradient tensor
The gravity-gradient tensor is given by
D(r0, t) = ∇δa(r0, t) = −(∇⊗∇)δψ(r0, t)(21)
where ’⊗’ denotes the Kronecker product (also known as dyadic or tensor product). It is a symmetric
tensor. Substituting equation (10) for the gravity potential, one obtains a tensor that can be divided
into four parts, distinguished by their dependence on directions. The ﬁrst part is proportional to the
14 J. Harms et al
local density perturbation:
D1(r0, t) = −4πG δρ(r0, t)er⊗er(22)
It is the only contribution with non-vanishing trace. Using Tr(a⊗b) = a·b, one obtains
Tr(D1(r0, t)) = −4πG δρ(r0, t),(23)
consistent with Poisson’s equation. The second part can be cast into the form
D2(r0, t) = −6G
S(θ, φ) = 5(ex·er)(ez·er)(31 −7er⊗er)
+ 4(ex⊗ez)sym + 5((ex×er)⊗(ez×er))sym.
Here (a⊗b)sym ≡a⊗b+b⊗a. The third part is given by
D3(r0, t) = 2G
0α26M0(t−r0/α) + r0
·(S(θ, φ)+(ex⊗ez)sym ),
and the last part is proportional to the delayed moment function
D4(r0, t) = −2G
Note that the unit vectors exand ezare not arbitrary coordinate axes, but correspond to the fault normal
and slip direction. The full gravity-gradient tensor is simply the sum of these four contributions. For
small times relevant to prompt perturbations, t<r0/α, the ﬁrst and last two contributions vanish
since M0(t)=0for t < 0, and the integral of the second contribution can be rewritten as
D(r0, t) = −6G
None of the four contributions vanishes for t→ ∞. Instead the time derivatives of the moment
function go to zero, and the moment function itself can be substituted by its ﬁnal value M0(t→ ∞).
The result is a gravity-gradient tensor whose components decrease with 1/r3
0, and is consistent with
the static gravity perturbation found by Okubo (1991) for shear dislocations in a half space, provided
that his result is evaluated for an event far from the surface.
4.3 Gravity-strain response
The instruments anticipated to observe weak gravity transients of the form given in equation (28)
measure changes in distance between test masses or relative rotations between suspended bars. These
Transient gravity perturbations induced by earthquake rupture 15
instruments measure the so-called gravity-strain tensor h≡I2[D](t), rather than the gravity-gradient
tensor D. The concept of gravity strain is similar to that of elastic strain. While elastic strain deter-
mines the relative change in distance between test masses bolted to an elastic medium, gravity strain
causes a relative change in distance between freely-falling test masses. The concept of gravity strain is
predominantly used in the gravitational-wave community, and since it emerged from studies of gravity
in the framework of general relativity, it should also be emphasized that some of the predicted effects
on instruments cannot be explained by Newtonian gravity theory. Nonetheless, in the context of this
paper, it is sufﬁcient to understand gravity strain as if produced by a tidal force and therefore it can be
calculated from the gravity-gradient tensor by double integration as stated above.
There are two types of gravity strain measurements: (1) the conventional scheme, which is a
measurement of the relative displacement of test masses typically carried out along two perpendicular
baselines (arms); and (2) measurement of the relative rotation between two suspended bars. In either
case, the response of a sensor is obtained by projecting the gravity strain tensor onto a combination of
two unit vectors, e1and e2, that characterize the orientation of the detector, such as the directions of
the two bars in a rotational gravity strainmeter, or the two arms of a conventional gravity strainmeter.
This requires us to deﬁne two different gravity strain projections. The projection for the rotational
strain measurement is given by
where the subscript ×indicates that the response to gravity ﬁelds is based on a relative rotation of the
bars. The vectors er
2are rotated counter-clockwise by 90◦with respect to e1and e2. In the
case of perpendicular bars er
2=−e1. The corresponding projection for the conventional
gravity strainmeter reads
The subscript +indicates that the response is based on a distance change between test masses. It is
straight-forward to evaluate the projections of the strain tensor in the form given in equations (22),
(24), (26) and (27) using the relation e>
Transformations between coordinate systems governed by a rotation matrix Rcan be applied to
the strain tensor in the usual way
0, t) = R>·h(r0, t)·R, (31)
and the projection is now carried out expressing the unit vectors in the transformed coordinates. This
transformation can also be interpreted as a real rotation of the detector frame representing a change in
detector orientation. For the case of perpendicular bars or arms of the gravity strainmeter, we calculate
16 J. Harms et al
the detector response to gravity strain before the arrival of P waves, as given in equation (28). Denoting
the square-root of the average of the squared strain amplitude over detector orientations (represented
by the frame rotations R), fault orientation, and slip direction by h·i, one obtains:
hh+(r0, t)i=hh×(r0, t)i=6p14/5G
where I4is the forth time integral. This expression allows us to make simple evaluations of expected
gravity strain signal amplitudes without having to specify directions and orientations. Adopting similar
assumptions as in section 4.1, we derive the following order-of-magnitude estimate of the amplitude
of prompt gravity strain signals:
hhi ∼ 6 104m5s−7t7
∼10−12 (t/10 s)7
As explained in the following section, there are at least three detector concepts that can reach these
4.4 Sensitivity of low-frequency gravity strainmeters
We conclude with a brief discussion about instruments and instrumental concepts potentially able
to detect the gravity transients from earthquakes discussed in this work. Several concepts have been
proposed for gravity strainmeters that target signals between 10 mHz – 10 Hz. These include atom-
interferometric, laser-interferometric, torsion-bar, and superconducting gravity strainmeters (Moody
et al. 2002; Harms et al. 2013). While none of the concepts has reached the sensitivity yet required
for the detection of earthquake transients, sensitivities of 10−15 Hz−1/2in the region 0.1 – 1Hz seem
within reach. Correspondingly and according to equation (33), transient gravity signals from earth-
quakes at ∼10 s after rupture onset and at epicentral distances ∼50 km can be expected to have
high signal-to-noise ratio. We point out that these low-frequency strainmeters are much smaller scale
(of order 1 m to 10 m) than the km-scale GW detectors LIGO and Virgo, which operate at higher fre-
quencies. In fact, some of the modern concepts of low-frequency gravity strainmeters evolved from
well-known gravity gradiometer technology. We also want to emphasize that the sensitivity of low-
frequency gravity strainmeters required for the detection of earthquake transients lies well below the
sensitivity required for GW detection at the same frequencies, which is about 10−19 Hz−1/2at 0.1 Hz
(Harms et al. 2013). So it is conceivable that these instruments can be either early prototypes of future
GW detectors, or instruments speciﬁcally built for geophysical observations.
Groups working on the realization of these new concepts rely on modern understanding of ter-
restrial gravity noise, seismic isolation, and thermal noise acquired over the past decade. For seismic-
noise suppression, each concept follows a different strategy. Atom-interferometric detectors use freely
Transient gravity perturbations induced by earthquake rupture 17
falling ultracold atom clouds, and seismic noise couples strongly suppressed into the system through
the laser interacting with the atoms (Baker & Thorpe 2012). Superconducting gravity strainmeters
achieve seismic-noise reduction by common-mode suppression by differential readout of test-mass
displacements versus a common rigid reference (Moody et al. 2002). Laser-interferometric concepts
rely on active and passive seismic isolation of suspended test masses (Harms et al. 2013). Torsion-bar
detectors proﬁt from the mechanical ﬁltering of seismic noise through a potentially very low-frequency
torsion resonance (M. Ando et al. 2001). In reality, isolation performance is impeded by cross-coupling
between degrees of freedom allowing seismic noise to couple into the strainmeter output through un-
wanted channels. The goal is to improve the mechanical designs to suppress these couplings.
As explained by Harms et al. (2013), terrestrial gravity noise is expected to pose sensitivity lim-
itations at a level 10−15 Hz−1/2at 0.1 Hz mainly due to atmospheric infrasound, and possibly also
seismic ﬁelds (depending on the instrument site). It was proposed to use microphone or seismome-
ter arrays to coherently cancel associated gravity ﬂuctuations from the strainmeter data. However, this
technique is mostly unexplored and therefore its performance difﬁcult to predict. Some experience has
been gained with the cancellation of atmospheric gravity noise in gravimeter data (Neumeyer 2010),
but gravity-noise cancellation in gravity gradiometers is not equivalent to cancellation in gravimeters.
The best strain sensitivity at 0.1 Hz so far has been demonstrated with superconducting grav-
ity strainmeters reaching 10−10 Hz−1/2(Moody et al. 2002). Torsion-bar sensitivities have surpassed
10−7Hz−1/2at 0.1 Hz (Shoda et al. 2014). Naturally, work on detector designs needs to be accom-
panied by careful selection of instrument sites, which can have a big impact on ambient seismic or
infrasound ﬁelds (Berger et al. 2004; Brown et al. 2014), and also on the associated gravity noise.
In this paper we addressed for the ﬁrst time the problem of modelling transient gravity perturbations
produced by an earthquake during fault rupture. Since gravity changes propagate essentially instanta-
neously in comparison with seismic waves, gravity perturbations are generated at any distance as soon
as the rupture starts. In particular, a prompt gravity perturbation is expected before the arrival of P
waves, and has been computed here by an original analytical model, validated with numerical simula-
tions. The model indicates that it would be challenging to observe prompt gravity signals within 10 s
of the rupture onset with current instruments, but that they may be within the reach of high-precision
gravity strainmeters currently under development.
This study sets building blocks to evaluate the detectability and information content of prompt
gravity perturbations. The analysis of a potential application to earthquake early warning will be re-
ported elsewhere, in which a few seconds of gravity signal would enable warning signiﬁcantly faster
18 J. Harms et al
than conventional systems based on seismic data. The detection of prompt gravity perturbations may
also open new directions in earthquake seismology, since it consists in a direct measurement of the
mass redistribution during fault rupture. How complementary is the information about the earthquake
source contained in the transient gravity ﬁeld and in seismic waves remains to be assessed.
Further analysis is needed to extend this study to cases where the Earth’s surface affects more
signiﬁcantly the gravity perturbation ﬁeld. Surface effects can be studied numerically, as was done
here, but an analytical model of gravity perturbations in the presence of a surface would provide
APPENDIX A: CALCULATING GRAVITY PERTURBATIONS USING SPHERICAL
A straight-forward and possibly more intuitive, but slightly more involved method to calculate gravity
perturbations is to directly integrate over the ﬁeld of density perturbations produced by a double-
couple source according to
δψ(r0, t) = −GZdVδρ(r, t)
The integrand can be expanded into its spherical multipoles. The integration over the radial coordinate
ris divided into two intervals: 0< r < r0and r0< r. Over the ﬁrst interval, one obtains the exterior
spherical multipole expansion:
δψext (r0, t) = G
dr r2ZdΩ δρ(r, t)Rm
are the regular and irregular solid spherical harmonics (in Racah’s normalization), respectively, Ym
the spherical harmonics, and dΩ = dφdθsin(θ). The corresponding expression for the interior spher-
ical multipole expansion is
δψint(r0, t) = G
dr r2ZdΩ δρ(r, t)Im
Transient gravity perturbations induced by earthquake rupture 19
Inserting the density perturbation of equation (3) into the exterior multipole expansion of the gravity
potential we have
δψext (r0, t) = G
dr r2∆(r, t)
·ZdΩ RP(θ, φ)Rm
The integral over angles can be carried out by expressing the radiation pattern of the density ﬁeld in
equation (4) in terms of the spherical harmonics Y1
RP(θ, φ)=2p2π/15 Y−1
and making use of the orthogonality relation
This leads to
δψext (r0, t) = G4π
dr r4∆(r, t)
2(θ0, φ0)∗ZdΩ RP(θ, φ)Ym
Further simpliﬁcation shows that the angular dependence is given by the quadrupole radiation pat-
tern RP(θ0, φ0). The integral over the radius can be simpliﬁed considerably by integration by parts.
An expression for the interior multipole expansion is obtained by a similar procedure. The solution
δψ(r0, t) = δψext(r0, t) + δψint(r0, t)for the gravity potential perturbation can then be written in
the form given in equation (10).
The form of the early gravity perturbation, equation (12), suggests that one can obtain this re-
sult from a simple effective source term. Due to the quadrupolar form of the radiation pattern (i.e. it
involves only spherical harmonics of degree 2), we can represent the source by a mass quadrupole
Qij =ZdV δρ(3xixj−δij r2)(A.9)
and zero monopole and dipole moments. Such a quadrupolar density distribution generates the follow-
ing exterior gravity potential:
δψ(r0, t) = −G
The integral in equation (A.9) can be solved again by expanding the integrand into its multipoles. For
early times, t<r0/α, the same result is obtained for the gravity perturbation with the only non-zero
20 J. Harms et al
components of the quadrupole moment tensor
Qxz(t) = Qzx(t) = 3I2[M0](t)(A.11)
Thereby, we have established an effective source model for prompt gravity perturbations generated by
APPENDIX B: SOURCE MODEL OF NUMERICAL SIMULATION
We consider a circular pulse-like rupture of ﬁnite duration with cosine slip rate function and constant
˙s(r, t) = 2δ
where Θ(·)denotes the unit step function. In words, the fault slip δat distance rfrom the hypocenter
builds up over a rise time T. The rupture propagates radially outwards with velocity vrup , until it
reaches the distance vrupτ. Therefore, τcan be regarded as the duration of the rupture, but the ﬁnal
fault slip occurs at time τ+T.
The moment function for a circular fault rupture can be calculated as
dr r ˙s(r, t0),(B.2)
where µis the shear modulus. Using the slip velocity model of equation (B.1), the moment function
is found to be
M0(t) = µ δ v2
−6πu + 4π3u3+ 3 sin(2πu), t ≤T
2π(−3 + π2(2 + 6(u−1)u)), T < t ≤τ
−2π3−3u+ 2π2((u−1)3−3uv2+ 2v3)
+ 3vcos(2π(u−v))−3 sin(2π(u−v)) τ < t ≤τ+T
12π3v2, τ +T < t
with u≡t/T , and v≡τ /T . After the rupture ends at t=τ+T, the moment function stays at a
constant value of µ δ π(vrup τ)2, and according to equation (18) the gravity acceleration grows as ∼t2
until P waves reach r0.
Transient gravity perturbations induced by earthquake rupture 21
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We thank Eric Cl´
e for discussions during the initial phase of this study and Mauricio Fuentes for
assistance with veriﬁcation and simpliﬁcation of analytical derivations. This work was supported by
NSF Grants PHY 0855313 and PHY 1205512 to UF. BFW acknowledges sabbatical support from the
e Paris Diderot and the CNRS through the APC, where part of this work was carried out.
JPA acknowledges support by a grant from the Gordon and Betty Moore Foundation to Caltech. We
acknowledge the ﬁnancial support from the UnivEarthS Labex program at Sorbonne Paris Cit´
10-LABX-0023 and ANR-11-IDEX-0005-02) and the ﬁnancial support of the Agence Nationale de la
Recherche through the grant ANR-14-CE03-0014-01.
SPECFEM3D is available at http://www.geodynamics.org/cig/software/specfem3d.