lecture notes on geophysical data processing and lecture notes on applied geophysics pdf free download
Introduction to Geophysics – Lecture Notes
March 23, 20151
1.1 What is geophysics?
Essentially, as the word suggests, geophysics is the application of method of physics to the study of
the Earth. The rocks does not diﬀer only by their macroscopic or microscopic properties studied
ﬁeld geologists or petrologists. They also diﬀer by their chemical and physical properties. Hence
as the rocks diﬀer according to their origin, structure, texture, etc. they also diﬀer by their density,
magnetisation, resistivity, etc. The bad news is that the physical properties do not always clearly
correlates with geological classiﬁcations and do not necessarily easily translates into the geological
terms. What does this mean?
Lets take the density as an example. We have
a rock sample and we have measured the value
of density to be 2.60 g/cm3. According to this
value we could assume that the rock sample could
be, e. g. a limestone, some shale, compact sand-
stone, rhyolite, phonolite, andesite, granite, pos-
sibly some kindof schist and many others. The
physical properties does not directly refer to the
geological classiﬁcation. This is a principal prob-
lem of geophysics, however, as we will see later,
there are ways to overcome this.
So, what we can conclude from this example?
The geophysics is a kind of proxy in our attempts
to study the geological structures. It does not
“talk” to us in geological terms and we have to in-
terpret obtained physical parameters in a ge olog-
ical sense. Moreover, the icentrifugal forcenter-
pretation is not unique as we have seen in our
example. The successful interpretation is based
on experiences of an interpreter and on the a pri-
ori knowledge of the geological environment stud-
ied. In the terms of our example – if we know
that we are working in the crystalline complex
we can mostly likely leave sedimentary rocks out
of our interpretation and we are left with rhyo-
lite, phonolite, andesite, granite or schist. And if
Figure 1.1: Possible inferences of structure
we study the geological sources a little bit more
at depth (Musset and Khan 2000).
we could also discriminate between the rock types
When we have discussed the essential and unavoidable drawback of geophysics it is time to
look on the bright side. Why do people bother with geophysics, what problems can it solve and
how it can help me in my particular problem? The advantage of geophysics is that it is able to
image hidden structures and features inaccessible to direct observation and inspection. That from
measurements on the surface we can deduce what is in the depth. Moreover, we can measure
on traverses or even make a grid and hence obtain a proﬁle view, map or even a 3d image of a
subsurface. Compare this with a geological mapping where we study the outcrops and, if we are
lucky, have also a few trenches or boreholes. There we have just a surface situation and we can
only guess how the surface structures continues to the depth.
Imagine the situation illustrated in the Figure 1.1. From the surface geological mapping we
would see a sediment ﬁlled valley with inclined strata on both sides. However, we have no clue
how it looks like in the depth. Four possibilities are sketched in the ﬁgure. The surface geological
mapping cannot give us any hint which of these is correct unless we would make a line of boreholes.
This, of course, would be both – time consuming and extremely expensive. How the geophysics
can help us here?
Lets assume that we have collected geophysical data on a traverse across the structure. The
methods used were gravimetry, magnetometry and a DC resistivity proﬁling. The gravimetry
distinguish rocks according to their densities, hence if we see an increase in gravity we can assume
rocks with increased densities – e.g. a maﬁc intrusion (Fig. 1.1c). Decreased densities of rocks
would also decrease the gravity readings we can assume presence of rocks with low densities – e.g.
a salt dome (Fig. 1.1d). The magnetometry indicates rocks with increased magnetisation – in case
of maﬁc intrusion (Fig. 1.1c) we would see an increase in magnetometry data. The fault zones
are usually connected with low resistivities. Hence if Fig. 1.1b or Fig. 1.1d would be the case we
would see a low resistivity zone over the fault. In such a manner the geophysics can add to the
surface geological prospection. In this particular case the geophysics provided us with the third
(desperately needed) dimension. The only necessary condition is that the target structure must
diﬀer in some of the physical properties from its surroundings (Tab. 1.1).2
The gravimetry detects tiny diﬀerences in the gravitational force. Since, according to the New-
ton’s law, the gravitational force depends on masses of the bodies, it allows us to diﬀerentiate
underground bodies and structures if their densities diﬀer.
2.1 Newton’s Law of Gravitation
Everyone is familiar with the Earth pull or attractive force. It causes thing s to fall and is also
responsible for a pretty hard work if we need to carry stones to build a house. The man who
discovered that every mass attracts another one was Sir Isaac Newton. In 1687 he formulated his
discovery into the equation (2.1) called the Newton’s Law of Gravitation:
where F denotes the gravitational force,
is the universal gravitational constant (6:673
10 N(m=kg) ), m are weights of attracting bodies and r is the distance between them. This
equation (2.1) enables us to calculate a gravitational force the Earth is pulling e.g. a rock on the
where M is the weight of the Earth, m is the weight of the rock and R is the diameter of
E r E
the Earth. We can see that it is inconvenient to use and measure the gravitational force, since it
depends on weights (masses) of both bodiesM andm . Dividing both sides of equation (2.2) by
m we get:
Since the force is computed asweightacceleration we can transform the equation (2.3) into:
deﬁning the acceleration caused by the Earth. The acceleration g is called the “acceleration
due to gravity” or “acceleration of gravity”. The value of g on the Earth surface is 9:80665 m=s
which is often simpliﬁed to 10 m=s . The unit of acceleration of gravity – 1cm/s – is also referred
to as galileo or Gal, in honour of Galileo Galilei, who was the ﬁrst who measured its value. The
modern gravimeters are capable of readings with the precision of 0.001mGal (0.01m=s ).6 Gravimetry
Figure 2.1: Topographic corrections (Musset and Khan 2000)
2.2 Gravity ﬁeld of the Earth and data reduction
Because the Earth is not a perfect homogeneous sphere, the gravitational acceleration is not con-
stant over the whole Earth’s surface. Its magnitude depends on ﬁve following factors: latitude,
elevation, topography of the surrounding terrain, earth tides and density variations in the subsur-
face. Within the geophysical prospection, we are interested in the last one, which is usually much
smaller then the latitude and altitude changes. The removal of unwanted components is often
referred to as reduction.
Latitude correction. The reason for the latitude correction is two-fold. First of all, it is caused
by the Earth’s centrifugal force being added to the gravitational force (vector sum). This
decreases the gravitational force with an increase of a radius of rotation. Hence the smallest
gravitational force is on the equator (maximal centrifugal force) and the largest is on the
pole. Second, the gravitational force is further aﬀected by the fact that the Earth is not
spherical but ellipsoidal. This further decreases the gravitational force on the equator. Both
of these eﬀects could be removed by the International Gravity Formula:
g = 978031:8(1 + 0:0053024 sin 0:0000059 sin 2) mGal (2.5)
It is clear that the centrifugal force changes only in the N–S direction, not in the W–E.
As we have seen from the Newton’s Law of Gravity – equation (2.1) – the gravity decreases
with the square of distance. Hence, if we lift the gravimeter from the surface (or any other
datum), the gravity will change. To be able to compare data measured in diﬀerent elevations
we have to reduce them to a common datum. This correction for the topographic eﬀects has
several steps. Their description follows.
Free-air correction. This is the ﬁrst step of reducing topography eﬀects. It simply corrects for
the change in the elevation of the gravity meter, considering only air (hence a free-air) being
between the meter and selected datum (leftmost part of the Figure 2.1). To get the change
in gravity acceleration with height we can diﬀerentiate the equation (2.4):
g M 2g
= mGal=m: (2.6)
R R R
Raising the gravity meter (e.g. extending its tripod) decreases the measured gravity values
by 0:3086 mGal=m. Hence to measure with an accuracy of 0.01mGal we have to measure
the elevation of the gravity meter with an accuracy of 3cm. The free-air correction varies
slightly with latitude:
Clearly, the sin and h parts are very small and could be neglected and we end up with
the above mentioned ratio.
Bouguer correction. The Bouguer correction removes from the data an eﬀect of rocks laying
between the measured point and reference datum (Fig. 2.1 in the centre). This eﬀect wasGravimetry 7
Figure 2.2: Gravimeter proﬁle across Sierra Madera, Pews County, Texas, illustrating the im-
portance of terrain corrections (Hammer 1939).
ignored during the free-air corrections. Hence we add a slab with an average density of
surrounding rocks – the Bouger correction:
= 0:04192 mGal=m; (2.7)
where is the density of the Bouguer slab.
The free-air and Bouguer correction is often combined into an elevation correction:
g g g
E FA B
= = (0:3086 0:04192) mGal=m: (2.8)
R R R
Terrain correction. The Bouguer correction assumed the slab to be inﬁnite in the horizontal
direction. This is not true, due to a topography and Earth curvature. The correcton for the
Earth curvature is used in a large scale surveys and we will leave it out now. The topography
correction, however, might be important (Fig. 2.1, right). The hill above the Bouguer slab
with its own gravity force pulls in the opposite direction than the Earth, therefore decreasing
the measured acceleration (Fig. 2.2). In a similar way, the valley also decreases the value,
because when computing the Bouguer correction we have already subtracted the Bouguer
slab (with a ﬂat surface) and did not account for the missing masses of the valley. Hence
the terrain correction is always added.
There are several methods of calculating terrain corrections. In any of these we need to know
the relief to certain distance from the station in detail. The common method is to divide
the surroundings of the gravity stations into zones, estimate average altitude in every zone
and compute the gravity eﬀect of the zones. Several versions of these zones were already
published (e.g. Hammer 1939). The easiest way is to print the zoning chart (Fig. 2.3) into
the transparent sheet and overlay it over the topographic map. Then the average altitude
in each zone and compartment is estimated and the diﬀerence between estimated value and8 Gravimetry
Figure 2.3: Zone chart for use in
evaluating terrain corrections at grav-
ity stations (Hammer 1939).
station elevation is computed and a gravity eﬀect of this compartment is found in the table
(Figs. 2.4 and 2.5). Finally, sum of eﬀects in all compartments and zones forms the terrain
correction for the current station.
It is clear that this method of computing terrain corrections is very tedious. Hence now usu-
ally the computer programs compute the corrections based from the DEM (digital elevation
model of the terrain).
Note that the eﬀect of the innermost zone – A – is not computed in the table. The reason
is that in such small surroundings the terrain should be ﬂat if possible. The gravity eﬀects
of irregularities in such close vicinity is very large, precise topography maps in scales large
enough are not common and dense precise measurements of relief would be inadequately
Tidal correction. The tidal correction accounts for the gravity eﬀect of Sun, Moon and large
planets. Modern gravity meters compute the tide eﬀects automatically. For the older in-
struments, one must compute the corrections by himself, e.g. according to the Longman
(1959), or consider the tides as a part of the drift of the instrument and remove it via a drift
Drift correction. This correction is intended to remove the changes caused by the instrument
itself. If the gravimeter would be at one place and take periodical readings, the readings
would not be the same. These are partly due to the creep of the measuring system of the
gravimeter, but partly also from the real variations – tidal distortion of the solid Earth,
changes of the ground water level, etc.
The drift is usually estimated from repeated readings on the base station. The measured
data are then interpolated, e.g. by a third order polynomial, and a corrections for proﬁle
readings are found.
2.3 Gravity meters
Basically, there are two main gravity types – the absolute and relative gravity meters. In geophys-
ical prospection solely the former are used. They measure relative gravity – the gravity changes,
not the value of gravity itself. If we want to measure the absolute value of gravity we must use
some point with already known gravity value and start our measurements there. Then, by adding
the measured diﬀerences, we know the absolute values of gravity on all our points. However, this
is needed only in large scale mapping, where we want to add our data to already existing grid.Gravimetry 11
Figure 2.6: Schematic sketch of the
astatic gravity meter. The zero-length
spring supports the mass M and keeps
it in balance in a selected gravity ﬁeldg.
Measurements are done by rotating the
dial, which raises or lowers the measur-
ing spring and provides additional force
Mg to return the mass to the stan-
dard position (Milsom 2011).
Thegravitymetersusedforthegeophysicalprospectionaretheastatic gravity meters (Fig.2.6),
where the mass is hold by the measuring spring. Elongation of this spring is proportional to the
gravity force pulling the mass. The older models of gravity meters used a dial to raise or lower the
measuring spring to place the mass to a standard position. In modern gravimeters this is done
From the sketch it is clear that the springs used in the meters must be extremely thin and
sensitive. There are two main types of the springs – steel springs in the LaCoste-Romberg gravity
meters and quartz springs in the others. Currently, there are two manufacturers of the prospection
gravity meters – the Scintrex with quartz springs and Burris, resembling the old LaCoste-Romberg
and using the steel springs.
To minimise eﬀects of thermal changes, the spring is in thermally insulated chamber (vacuum
chamber), the new models are also equiooed with an additional heater to keep the internal tem-
perature as stable as possible. Hence, removing batteries from the gravity meter leads to change in
the inner temperature resulting in unstable and unreliable readings due to changes of mechanical
properties of the measuring spring. When the gravimeter is without the power supply for a long
time (several hours) it could take as long as 48hours before realiable readings could be done again.
2.4 Rock densities
The densities of rocks (Tab. 1.1), naturally, depends on the mineral composition of particular rock.
However, not only mineral composition, but also other factors inﬂuences densities. The next main
factor, mainly when dealing with sedimentary rocks, is the porosity and kind of media ﬁlling the
pores. Increasing porosity decreases the density, since air (or any other media ﬁlling the pores
(water, gas, oil, etc.) has lower density than any of minerals. The other factors are weathering of
rocks, fractures and joints, etc. Combining all these factors clearly explains the high variance of
measured values reported in a literature.
Densities needed for a data and interpretation interpretation (e.g. for the Bouguer anomalies)
could be either measured in laboratory (keeping in mind the laboratory values could diﬀer from
the reality mainly due to change of the media ﬁlling the pores in the nature and in laboratory) or
could be estimated from the gravity measurements.
One of widely used methods for density estimation is the Nettleton’s method (Nettleton 1939).
This method is based on the fact that the Bouguer anomaly depends on the density of rocks as well
as on the topography. If the topography along the proﬁle is changing but the density is constant
then, according to the equation (2.7), the Bouguer anomaly should be constant as well. If it is not
constant then the density estimate is wrong and the topography changes are not compensated well.
Therefore, if we compute a set of Bouguer anomaly curves with diﬀerent densities and compare
them with the topography the Bouguer curve which correlates the least with a topography is the
curve with a correct density estimate.12 Gravimetry
Figure 2.7: Illustration of the Net-
tleton’s method for density estimate.
The Bouguer curve which correlates
least with the topography is calcu-
lated with a correct density estimate.
The top graph shows the topography
along the proﬁle. In the middle graph
Bouguer curves for diﬀerent densi-
ties are plotted, the curve for the se-
lected density estimate is plotted in
red. The bottom graph depicts cor-
relation (Spearman’s correlation co-
eﬃcient) of topography and Bouguer
curves for diﬀerent densitites.
This method could be applied if a) the proﬁle crosses some distinct topographical feature, b)
the density of the subsurface is not expected to change substantially.
2.5 Gravity eﬀect of selected bodies
The simple geometrical bodies are often used for the modelling before the survey is carried out.
The aim is to get a rough estimate of the anomaly eﬀect of the target structure. If we, for example,
ﬁnd that the amplitude of a modelled anomaly is lower than sensitivity of our instrument then
there is no need to do any measurements at all...Estimating the amplitude and width of the
anomaly also enables us to plan a density of proﬁles and station spacing. For an interpretation at
least three stations within the anomaly are necessary.
There is a small number of simple basic bodies, however, combining them together can build
up even a complex model.
Gravity eﬀect of a sphere A sphere is the most basic body and usually is used as a part of
other models or could approximate symmetrical bodies. The gravity eﬀect of a sphere at
point P (Fig. 2.8) is:
g =g cos = = ; (2.9)
(x +z )
where is a density of the sphere,a is a radius of the sphere andz is the depth of the center
of the sphere.Gravimetry 13
Figure 2.8: Gravity eﬀect of a sphere (Telford et al. 1990)
The depth z of the sphere could be estimated from the measured anomaly. When g =
thenz = 1:3x . In other words, the depth of the sphere center could be estimated from the
half-width of the anomaly at half of its value (see Fig. 2.8).
Gravity eﬀect of a vertical cylinder The gravity eﬀect on the axis aces of a vertical cylinder
2 2 2 2 2
g = 2
L + (z +R ) (z +L) +R ; (2.10)
where L is the vertical size (length) of the cylinder z is the depth of its top and R is its
diameter. IfR1, we have an inﬁnite horizontal slab, which we used also for the Bouguer
g = 2
Gravity eﬀect of a horizontal rod Assuming a horizontal rod perpendicular to the x axis at
a depth z, the gravity eﬀect is:
m 1 1
g = ; (2.12)
2 1 1
h i h i
z 1 + 2 2 2 2
2 : ;
x +z x +z
1 + 1 +
(y+L) (y L)14 Gravimetry
Figure 2.9: Gravity eﬀect of a horizontal rod. a) Three dimensional view. b) Projection of the
plane containing the rod and the y axis. c) Projection on the xz plane. d) Gravity proﬁle along
the x axis. (Telford et al. 1990)
where m is the mass of the rod. If the rod is expanded into the cylinder with a dimension
a then m = a . When the length L of the rod is inﬁnite (usually a good approximation
when the L 10z then the (2.12) simpliﬁes into the:
g = : (2.13)
z 1 +
The depth z of the rod could be estimated from the half-width of the anomaly: z =x1.
A lot of other geometric bodies could be found in the literature, enabling us to build a complex
models. Examples of changes in gravity eﬀect of doﬀerent bpdies with depth could be found in
Figs. 2.10 and 2.11. Two examples of models using the spheres and cylinders are plotted in the
Figs. 2.12 and 2.13.
More complex modelling could be done using the computer modelling and irregular bodies.
Check, e.g. the Gordon Cooper’s web page at the University of Witwatersrand (Cooper 2012).Gravimetry 15
Figure 2.10: Anomalies of a sphere and a horizontal cylinder at diﬀerent depths. (Musset and
Figure 2.11: Anomalies of narrow sheets at diﬀerent depths and dips. (Musset and Khan 2000)16 Gravimetry
2.6 Gravity data processing
Once the gravity data are measured the more demanding task is to be carried out – the data
processing and interpretation. Some procedures of data processing has been already mentioned in
the section 2.2.
Figure 2.12: Approximation of an anti- Figure 2.13: Approximation of a salt di-
cline (Mareš and Tvrdý 1984). a) A geolog- apir (Mareš and Tvrdý 1984). a) A geo-
ical section of an anticline. b) A schematic logical section of a diapir. b) An approx-
representation with density distribution. c) imation of the diapir by two spheres. c)
An approximation of the anticline by three An approximation of the diapir by a verti-
inﬁnite horizontal cylinders. cal cylinder.
The ﬁrst step in the data processing is deleting of wrong gravity readings. During the the ﬁeld
measurements there is usually several gravity readings taken at every station. Now, the outliers
are removed and the rest of gravity readings from every station are averaged.
Next, the readings from the base station are taken to determine the drift of the instrument.
First, these data need to be corrected for the diﬀerent heights of the tripod, the free-air correction
– equation (2.6). Second, the drift should be estimated – usually the data are interpolated using
the second or third-order polynomial (Fig. 2.14).
Third, the readings at individual stations are corrected from the drift. The drift is estimated
from the ﬁtted polynomial according to the time of the gravity reading.Gravimetry 17
Figure 2.14: Drift correction. The blue line shows the gravity readings at the base station
corrected for the free-air. The red line depicts the third-order polynomial ﬁtted into the gravity
readings. This polynomial will be used to estimate the drift values for the readings at individual
Fourth, the drift corrected data are reduced again, now using the latitude, free-air and Bouguer
reductions – see equation 2.8). If necessary, the density for the Bouguer slab is estimated (e.g.
There are also additional steps, which depends on the type of the survey and target structures.
However, usually we want to suppress regional anomalies and enhance the local ones or vice versa.
The regional anomalies is a general term depending on the size of target structures. These
anomalies are caused by large and deep structures, often larger than our survey area. In the data
they usually represent the long-wavelength high-amplitude anomalies (Fig 2.15c, d). Sometimes
they are also referred to as a trend (Fig. 2.15d). There are numerous techniques to remove the
trend, the easiest are based on approximation by a polynomial. In this case we take the part
of the data without our target anomaly and ﬁt a polynomial through them. This polynomial
approximates the eﬀect of large-scale regional structures and we can subtract it from our data
leaving us with residual anomalies. The residual anomalies are, in an ideal world, anomalies
caused only by our target structures.
2.7 Gravity data interpretation
The interpretation of gravity data could be only a simple qualitative analysis in a way: “Look,
there is a sharp local decrease of gravity, this could be a cave” Or a more complex quantitative
analysis, where, based on the qualitative assignment, we try to somehow model the subsurface.
In this respect we have to bear in mind that the interpretation (inversion) of geophysical data is
non-unique. In gravity prospection not only that diﬀerent bodies could have similar anomalies
(cf. ﬁgures 2.8 and 2.9) they can also produce exactly the same anomaly (Fig. 2.16). The non-
uniqueness is inherent to gravity data and could not be overcome e.g. by adding more gravity
data. The only way how to get sound and reliable interpretation is to include an a priori geological
knowledge and, if possible, also data from another geophysical methods.
between the bodies. Hence a sphere with the density of 2:3 g=cm surrounded by the rocks with a
density 2:5 g=cm will produce exactly the same anomaly as a sphere with the density of 2:5 g=cm
surrounded by the rocks with a density 2:7 g=cm . In a similar manner a half slab on one side of a
fault with a positive density contrast could produce the same anomaly as a half slab on the other
side with a negative density contrast (Fig. 2.17).18 Gravimetry
Figure 2.15: Illustration of regional and
residual anomalies (Musset and Khan 2000).
The observed gravity curve contains informa-
tion about all geological structures (topmost
curve). If we are looking for the dyke, then the
anomalies due to the dipping strata and granitic
pluton are not relevant to our research and we
would like to remove them from the data to ease
Figure 2.16: Non-uniqueness of the gravity interpretation. The plotted models produce exactly
the same gravity anomalies (Musset and Khan 200).Gravimetry 19
Figure 2.17: Two half slabs with the same
anomaly (Musset and Khan 2000).
Figure 2.18: Depth rules for various bodies (Musset and Khan 2000).20 Gravimetry
As we have already seen in the section 2.5,
simple rules for estimation of depth of some sim-
ple bodies could be derived. There are some more
in the ﬁgure 2.18. These rules are useful for esti-
mating initial parameters for further modelling.
Building models from simple geometrical bod-
ies is easy, however, one can easily see that not
all the geological bodies could be easily approx-
imated by them. Therefore, another modelling
techniques were developed.
One of these is building the models from poly-
gonal bodies with arbitrary number of corners
(Figs. 2.16b, 2.19). There are formulas for com-
puting gravity eﬀects of such bodies, often based
on the original Talwani’s algorithm developed in
late ﬁfties (Talwani 1959). The modelled bod-
ies are usually 2D (inﬁnite in the y direction) or
2.5D (bodies have limited length in the y direc-
tion). However, equations for 3D modelling are
also available. The advantage of this (polygonal)
approach is in fact that the computations are fast
and memory cheap and can be easily run on any
of current computers. There are number of such
computer programs available here and there (e.g.
Cooper 2012, Fig. 2.19). The usual approach is
to build an initial model estimated according to
Figure 2.19: Modelling a maar-diatreme
the measured gravity data and geological evidence
volcano as a set of 2.5D bodies. In the top
and then using either a trial and error technique
graph the measured gravity curve (broken line)
and to some extent also with automated inver-
and computed curve for the current model
sion procedures we try to match the measured
(solid line) can be seen.
and modelled gravity curves. Due to the non-
uniqueness of the gravity data, there are, unlike
to some other geophysical methods, no “black-box” automated inversions. Currently, it is not
possible to put data into some computer program, press a button and get the result. Although
there are some attempts to achieve this.
Another approach is to divide the model into regular cells and assign each cell some density
value (e.g. Snopek and Casten 2006). The cells are usually cubes, but any geometrical repre-
sentation is possible. This kind of discretisation is common to many geophysical methods. The
advantage is that arbitrarily complex models could be achieved. The drawback – the number
of cells (and hence also the parameters) very quickly arises, mainly for the 3D case, and the
computing gravity eﬀect of such model is (computationally) very expensive.Gravimetry 21
2.8 Applications of the gravity method
The gravity measurements could, obvi-
ously, be applied anywhere where suﬃcient
density contrast is expected. Neverthe-
less, there are situations and ﬁeld condi-
tions suitable and unsuitable. Lets start
with the former.
One of the scenes where gravity excels
is a regional geological mapping. It is due
to the fact that the gravity meter is easily
portable, does not need any wires and ca-
bles and one or two people are enough to
operate it. Therefore, there is no logistic
problem in measuring long (several kilome-
tres or even more) traverses. Another ad-
vantage is the high depth reach – it is com-
mon to model structures in the depth of
several kilometres (Fig. 2.20, 2.21).
Another ﬁeld where the gravity mea-
surements are indispensable is the mapping
physical methods that could directly detect
voids (ground penetration radar being the
second). Therefore, the gravity method is
Figure 2.20: Bouguer anomaly map of the NW
often used to search caves (Fig. 2.22), old
Czech Republic (Blecha 2009).
mines and galleries, or diﬀerent voids and
cavities beneath the roads.
Very common application of the gravity method is mapping of the sedimentary basins for the
oil prospection. If the densitites of sediments are known (e.g. from boreholes) then not only the
lateral extent but also the depth of the basins could be mapped. Another example from the oil
prospection is the mapping of salt domes, since they often form oil and gas deposits.
This set of examples could be ﬁnished by a volcanological example – maar-diatreme volcanoes
are often mapped using the gravity data. The eruption of the maar volcano is very forceful, the
explosion creates a large crater and shatters the country rock. After the explosion, part of the
material falls back to the crater, however, is "ﬂuﬀed up" by the explosion and hence its density is
(Fig. 2.21). However, the best results are always obtained by combination of several geophysical
methods. An example combining gravity and resistivity data is plotted in the ﬁgure 2.24.22 Gravimetry
Figure 2.21: Geological model along NW-SE gravity proﬁle. Rock densities D are in kg/m 3. Kr
- Krudum massif, intrusion of YIC granites. Inlet a): the depth of pluton ﬂoor would increase by
about 2 km in case that the high density lower part of Saxothuringian nappes is omitted. Earthquake
hypocenters less than 8 km from the proﬁle are indicated by black circles. The upper panel shows
the plan view of the model at a depth of 2 km b.s.l.; crossed areas in amphibolites (green color)
indicate two diﬀerent types of rocks in the area of plan view. (Blecha 2009).Gravimetry 23
Figure 2.22: Gravimetry proﬁles showing possible occurrence of an underground cavity.