Publication n°121 of the International Association of Hydrological Sciences Proceedings of the Anaheim Symposium, December 1976

GEOPHYSICAL METHODS FOR LOCATION OF VOIDS AND CAVES

Roy J. Greenfield, Peter M. Lavin, and Richard R. Parizek Dept. of Geosciences, The Pennsylvania State University University Park, Pennsylvania

Abstract The application of microgravimetric surveying, subsurface temperature

measurements, electrical resistivity surveying, and seismic surface waves to the location of underground voids and caves will be discussed. Examples of field studies in carbonate terrains of central Pennsylvania, U.S.A., are presented. Microgravimetric procedures for determining soil thickness variations related to subsurface cavities as well as direct detection of cavities will be described. The existence of a cave may be inferred from subsurface temperature measurements at depths of 5 to 10 ft. Electrical resistivity measurements over fracture traces that delineate underlying zones of nearly vertical fracture concentration and caves show that these features give an anomaly. However, the anomaly can be either that for a resistive or conductive zone, depending on the time since a substantial rainfall. Theoretical results are presented for a resistive horizontal cylinder representing a dry cave. Use of seismic surface waves generated by harmonic or transient sources is discussed. Both a field study in an area of fractures and an ultrasonic seismic model study of a shallow cave show these features are detectable through changes in amplitudes and frequency content.

Introduction For a number of years graduate students and staff at The Pennsylvania

State University have been studying the geology and hydrology of carbonate terrains, with particular emphasis in central Pennsylvania. In the course of this work a number of studies have been conducted using geophysical methods to locate and to determine the extent of zones of fracture concen­tration revealed by fracture traces and lineaments, to locate caves, and to define the depth and configuration of the weathered top of bedrock below residual and transported soils. This paper contains a brief review of pertinent results of only some of these investigations using gravity, subsurface temperature, electrical resistivity, and seismic techniques.

Significance and Nature of the Problem Detailed soils, subsurface geological and hydrological data are re­

quired to resolve land use planning, foundation engineering, geotechnical and environmental pollution control problems that arise within carbonate-terrains in all regions. Test drilling is costly and there is a need to extrapolate results between boreholes and to define and isolate discontin­uities not detected by test drilling. Carbonate terrains frequently dis­play complex and rather abruptly varying subsurface conditions that are important in all kinds of foundation and resource exploration, evaluation and design work (see Parizek, 1977, in this volume). Of particular con­cern to some are the depth to and configuration of the bedrock surface. Weathering of carbonate bedrock is facilitated by joints, fractures, bed­ding plane partings, individual beds, faults, zones of fracture concentra­tion, and related structures that are dispersed and that intersect in a difficult to predict manner (Parizek, 1977, in this volume).

Another class of problem includes the detection of sediment-filled

465

and/or open voids in bedrock. These may be interconnected regionally and serve as major conduits to groundwater and pollutants, or they may be abandoned and located well above the seasonally high water table elevation.

Geophysical investigations can be helpful in addressing many of these data acquisition problems. New methods still need to be perfected to improve on the size and depth of features that can be detected which varies in importance depending on the problem at hand. This paper reports on four geophysical methods that have been found useful under Pennsylvania's geo­logical conditions. Our research is continuing and we realize that there are limitations to our findings. Survey procedures applied and data re­quirements of others will not always be met using these techniques because of local geological difference and limitations of the methods.

Microgravimetric Methods The success of gravimetric techniques in mapping near surface density

changes due to buried stream valleys, bedrock topography, fracture zones and cavities (natural and man-made) is well established. Excellent re­views of procedures and results using standard gravimeters have been given by Neumann (1967) and Arzi (1975). A review of the application of "micro-gal" gravimeters to engineering problems was presented at the annual Meeting of the Transportation Research Board in 1975 by Omnes (1975). To review these "reviews" would do a disservice to the authors; therefore only certain aspects of microgravimetric methods that supplement these papers will be presented. No consideration will be given here to problems that are solvable using normal surveying procedures to detect anomalies on the order of 0.2 mGal or more (with station spacings of 300 ft or more) due to large caves (e.g., Colley, 1963) or buried stream valleys (e.g., Rankin and Lavin, 1970). Rather, we will concentrate on so-called high accuracy or microgravity surveying. Even this area requires breaking the discussion into two parts, deDendent on whether standard gravimeters or "microgal" meters are employed.

Neumann (1967) and others have discussed the procedures and require­ments for meaningful results using standard meters. The required accuracy of the surveys is dictated by the size of the expected anomalies. Numer­ous model studies have been made to study this aspect of the problem. Figure 1 is an example, showing the size of the anomaly as a function of the depth and size of the source for a sphere model.

Standard gravimeters have a reading precision of about 10 yGal; there­fore, extreme care must be exercised in reading the instrument and correct­ing for its drift characteristics. Multiple readings at a given station and reoccupation of a single base station at approximately one-half hour intervals are often necessary. Corrections for elevation and position require careful (but easily attainable) surveys and should not produce significant errors with the possible exception in areas of considerable topographic relief where the choice of density to be used for the eleva­tion and terrain corrections may be difficult. Study of the correlation between gravity anomalies and topography forms an important part of any analysis under these conditions. Neumann (1967) considers that it is quite feasible to detect anomalies as small as 30 yGals when sufficient care is exercised and station spacings are small enough so that the ano­maly is defined by values at several stations.

Typically, the noise level of high accuracy surveys is high, often reaching levels on the order of the desired anomalies. Most of the sources of noise (such as near surface lateral density variations, changes in depth to bedrock, man-made structures) are inescapable. Thus, data enhancement techniques (such as downward continuation and vertical derivatives) are

466

MAXIMUM GRAVITY ANOMALY FOR A SPHERE OF DENSITY CONTRAST 2.5 gm/cc

100

LU H Z LU O

Q. LU O

10 20 RADIUS (ft)

Figure 1

30

rarely called for. (In fact, one often has to low pass filter the data prior to attempting a quantitative interpretation, a point which will be demonstrated shortly.) Nevertheless, Fajklewicz (1976) proposes using measurements of the vertical gradient of gravity (over a distance of 1 yd) for detecting underground cavities and tunnels resulting from mining. The major effect of topography and the various sources of geologic and man-made noise on gradient measurements compared with that on gravity val­ues limits the application of the gradient techniques to areas of small topographic variation and relatively homogeneous conditions.

Normal interpretation involves curve fitting procedures, usually using simple geometric models such as spheres, hemispheres, cylinders and slabs. Numerous examples are presented in the literature previously mentioned (and the many references cited in those works). A somewhat dif­ferent approach will be presented here in order to illustrate the problems caused by "geologic noise."

A detailed gravity survey was run by Galiette (1974) in selected areas where well or seismic depth control was available as part of a study to determine soil thickness and bedrock competency for a sewage effluent spray irrigation program. The bedrock in the test area consisted of massive do­lomite with interbedded sandstones. Variations in relief due to solution cavities and fracturing characterize the bedrock surface. The soil consists of thin layers of clay, silt and sand. An intermediate mixed zone of soil and large fragments of bedrock is common. Even though the area is complex, it was assumed that it could be modeled by a two-layer system with a density contrast of 1.0 g/cc (based on measured values).

Gravity readings were obtained at 20 ft intervals along selected pro-

467

files. A standard Worden gravimeter was used. Drift correction was based on reoccupation of the base station at 30 min intervals. Relief in the area was moderate (20 ft) so no terrain corrections were applied. The data were reproducible within ±0.05 mGal and the simple Bouguer anomaly was judged to be good to ±0.07 mGal.

The data were interpreted using a modified version of a method pre­sented by Parasnis (1966) for the purpose of determining the minimum value of the maximum possible relief on the bedrock surface in a given region. In effect, the method is the equivalent stratum technique without downward continuation of the gravity data. The equivalent stratum method (Grant and West, 1965) relates the topographic relief h(x,y) on a surface at a mean depth, h, to the observed gravity anomaly continued down to the mean surface, Ag(x,y,h), by

Ag(x,y,h) 2irGAp

h(x,y) =

where G is the gravitational constant and Ap=p -p is the density contrast across the surface (Figure 2). The procedure requires that the mean depth to the surface (as well as the density contrast) be known and that the relief be small compared to this mean depth.

Because of the high noise level of the data relative to the desired anomalies and the fact that the depth to bedrock (continuation distance) was only a few tens of feet, the gravity data were not continued downward prior to determining the variation of the bedrock surface. If the gravity difference (Ag) at the surface between two stations is used, it can be dem­onstrated that the relief (h) on the bedrock surface between the stations

Ag h >

27TGAp

Thus, the use of the values of the gravity anomaly between successive grav­ity stations gives a measure of the minimum relief on the bedrock between those stations. It requires that the depth to bedrock be known at one point in the survey area and that Ap be known. Since the gravity anomaly is not downward continued, it is smoother and lower in magnitude than it would be if the equivalent stratum technique was used. Thus, the estimate

MODEL FOR BEDROCK RELIEF

STATION 1 STATION 2 A 7C DATUM

ROCK Figure 2

468

(0

O 0

0.1

0.2

0.3

BEDROCK SURFACE DETERMINED FROM FILTERED GRAVITY ANOMALY

S N 200 400 600 f t

£ 20

Z

1 ' UJ

ûj - 2 0

FILTERED ANOMALY

BOUGUER ANOMALY

WELL CONTROL

ROCK

Figure 3

of the variation of h is smoother and smaller than the actual surface. This was not found to be a problem as long as the maximum depth to bedrock in the area was less than 100 ft.

Figure 3 shows the results from the work of Galiette (1974) for a typical profile in the test area. The Bouguer anomaly and filtered anomaly (using a 5 point operator with a cutoff of 0.005 cycles/ft) are shown in the upper graph. Use of the unfiltered data generally leads to an unstable solution as many of the short wavelength features (< 50 ft) are probably not related to the bedrock surface; they represent the "geologic noise" in this area. Using the known depth at the well, the values of gravity were converted into depth values (lower graph). The bedrock surface shows a depression (possibly a zone of concentrated weathered bedrock) with a lateral extent of 140 ft and a depth of 26 ft. There is a slight depres­sion on the topographic surface directly above this feature. The calcula­ted depth on the north end of the profile agrees closely with the depth at a nearby test well. Calculation of the theoretical gravity response due to the computed bedrock surface (using a two-dimensional polygon model) agrees well with the filtered anomaly.

In general, the interpretational method can be expected to give a smoother and shallower bedrock surface than is actually present. In the test area, depth values were no more than 20% too shallow as long as the bedrock interface was within 100 ft of the surface and the control depth was no greater than 50 ft. Choice of a density contrast (if not known) determines to some extent whether a minimum or maximum estimate of depth is obtained.

Turning now to the use of "microgal meters", it must be remembered that the limiting factor in gravity detection is not the sensitivity of the instrument (although problems of unstable operation are possible if

469

the instrument is too sensitive). Rather, the limiting factor is the accuracy with which the corrections can be made and the degree of varia­tion in local geology, topography and man-made sources of noise. In areas of small topographic relief with good maps and density control, anomalies as small as 30 yGal may be significant. In areas of rough topography anomalies would probably have to exceed several tenths of a milligal to be considered significant.

The reading precision of the microgal meters is about 2 to 5 yGal. Used with care, readings may be obtained faster than with standard meters as the need for repeating readings to achieve the desired accuracy is reduced (or eliminated).

The newsletter of the Franklin Institute in Philadelphia of November, 1959, reports on the use of a conventional Worden gravimeter, modified to have a sensitivity of about 4 yGal, to detect cavities under the city streets (Glenn, 1959). No corrections to the data were required as the effect of the cavities was very sharp compared with the "gradual" varia­tions due to changes in elevation, position and "topography" (buildings). Cavities as small as 23 ft3 were found.

A similar meter was tested and used by one of the authors (PML) in 1960. Figure 4 shows the drift-corrected readings over two known manholes in a street. The gradient is due primarily to the change in elevation along the profile. The absence of a major connection under the street between the manholes is indicated by the relative gravity high in the cen­ter of the profile. Further use of the instrument to search for small solution cavities (- 60 ft ) immediately below an airport runway met with limited success as the "noise" (due, in part, to variations in the thickness of the macadam and base courses) resulted in anomalies similar to those expected from the cavities.

GRAVITY ANOMALY OVER MANHOLES

Figure 4

470

TEMPERATURE PROFILES

AT BOTTOM OF 6 FT HOLES

DISTANCE (FT) S N

80 60 40 20 O 20 40 60 ~T~

Figure 5

Starting in 1968, Compagnie Générale de Géophysique used a specially designed "Microgal" gravimeter (manufactured by LaCoste and Romberg, Inc.) for detection of cavities (Omnes, 1975). The reported "measuring accu­racy" of about 5 yGal and the reduction of error by a factor of 5 in the results when compared with those obtained from a standard gravimeter led to a marked increase in the use of the tool to the point where it became their sole method used for detecting cavities for highway projects (Omnes, 1975). Some of the success of the gravity method is undoubtedly due to the fact that the observed anomalies are always larger than the theoreti­cal anomalies based on the actual dimensions of the voids. This is due to the effect of the cavity on its surroundings (in terms of stress relief and increased fracturing and dissolutions). Several examples of this phenomenon are presented in the review by Omnes (1975).

Subsurface Temperature Measurements Shallow subsurface temperature measurements have been used in the past

to determine the location of aquifers (e.g. O'Brien, 1970; Cartwright, 1968; Bair, 1976). O'Brien (1970) used time averaged values to determine

471

the spatial distribution of aquifer properties west of Schenectady, N.Y. Here the cool aquifer depressed the time average (mean) of the temperature at a 6 ft depth by several degrees. The spatial variation of the decrease of the average temperature was found to be an indicator of water transmis-sivity variations.

Work was done by Ebaugh (1973) and Ebaugh, Parizek and Greenfield (1974) to determine if soil temperature measurements could be used to locate caves or zones of fracture concentration in the folded and faulted carbonate areas in central Pennsylvania. Three sites were studied; we will discuss one of these, the Miller Cave site. Miller Cave is 1 to 10 ft wide and 15 ft high on the average. Its top is 30 ft below the surface. Temperature profiles at 6 ft depth taken perpendicular to the cave are shown in Figure 5. Note the higher 6 ft depth temperatures over the cave

Annual air temperature wave

Q lowrt, wet

Soil temperature pro

- ' 1 7 ° F ,—— — J18.8°F,—"—"" J-n"F, —.-

Secondary permeability (Joints, zones of f racture concentrat ion, faults, bedding planes)

HEAT FLOW MECHANISMS IN MOIST SOIL Over fracture Oft fracture

CONDUCTION Moist soil, high thermal diffusivity, «

Wet soil, low then diffusivity <>

CONVECTION Laiger heat movement by mass transfer

Smaller heat movement by mass transler

Q

THEORETICAL ANNUAL SOIL TEMPERATURE VARIATION

ABOVE AND ADJACENT TO FRACTURED BEDROCK

Figure 6

472

in the summer (e.g. 8-21-72 reading) compared to the summer temperature away from the cave. Also note that the winter temperatures are lower over the cave than the winter temperature away from the cave. Thus the measure­ments indicate a larger peak to peak amplitude for the annual temperature wave component over the cave (22°F) than at the portions of the profiles away from the cave (18°F). This larger amplitude may reflect a higher thermal diffusivity of the soil above the cave. Thermal diffusivity is dependent on the fractional volume of pore space filled by water; the value is maximum at approximately a 15% volume fraction (Bauer et al., 1972). Thus, if the cave and overlying zone of fracture concentration allow the soil above the cave to drain (moist soil) the thermal diffusivity would be higher than in areas with poorer drainage (wet soil). This leads to the greater observed amplitude for the annual component of 6 ft temper­ature fluctuations above the cave. This effect is shown diagramatically in Figure 6. The indication of drainage of the soil above Miller Cave is con­sistent with the drainage inferred from electrical resistivity measurements (see section on Electrical Resistivity). Figure 6 also shows how the added water flowing into the fracture above the cave could carry heat (by convec­tion) downward in summer, and convect cold surface water in winter.

The question of whether the temperature in the cave could directly affect the 6 ft deep temperature values was examined theoretically. The effect being considered is similar to the lowering of the mean temperature, at 6 ft, by the cool aquifer noted by O'Brien (1970). In O'Brien's work the lateral extent of the aquifer was great enough so the heat flow was essentially vertical. However, it is more realistic to model a cave as an infinitely long horizontal cylinder, at a constant temperature which is different from the mean annual air temperature. The solution for the

Figure 7

473

steady state temperature distribution for this cylinder model in a half-space (Figure 7) was evaluated by Ebaugh (1973). The solution was adapted, by analogy, from the electrostatic solution given by Smythe (1950, p. 76). The calculation gave the distribution of temperature about a cave whose mean annual temperature is different by two degrees from the mean annual air temperature, as is the case in the study area. Figure 8 shows the results for a model cave representative of the Miller Cave geometry. The theory predicts an anomaly, in the mean temperature, at a point directly above the cave, of 0.13°F at 3 ft depth and of 0.25°F at 6 ft depth. These values are too low to explain the measured anomaly. In general a cave will not give a measurable temperature anomaly by conduction of heat to the cave. As noted, however, a cave may be located by its effect on drainage of the overlying soil or, as pointed out by Ebaugh, Parizek and Greenfield (1974), by temperature variations noted in deeper boreholes or boreholes that pass near the cave.

Electrical Resistivity Electrical resistivity surveys are used to define the subsurface re­

sistivity structure. A general review of these methods is given by Van Nostrand and Cook (1966).

The most obvious application of electrical resistivity methods to subsidence monitoring is the location of caves or voids, since a cave or void has effectively infinite resistance unless filled with water or soil. The difficulty is that the cave will not be detectable if the depth to its top is greater than approximately one-fourth its height. Figure 9 shows the geometry for a Wenner array traverse over a horizontal cylinder model for a cave at four tenths of the cave radius below the surface. The appar­ent resistivity profile for this model is shown in Figure 10 for different ratios of "a" spacing to size of the cave (Yu, personal communication).

TEMPERATURE ANOMALY DUE TO CAVE, THEORETICAL DISTANCE (FT)

0 20 40 60 80

u. o

g 0.4 3 H

<r 0.6 UJ 0.

2 UJ 0.8

1.0

- 3FT____„—-—-——~~~~^Z-^——~Z^~——- ~~—"

6FT^_~—""""^ ^^-~^~^ ^^-—

- 10FT^^-^ s ^

~ 20FT/ DIAMETER OF CAVE Y DEPTH TO CAVE'S AXIS

> ^ TEMPERATURE DIFFERENCE

a

100

15 FT 37.5 FT 2°F

Figure 8

474

WENNER ARRAY ELECTRICAL RESISTIVITY OVER A CAVE

.4 R(

V^a

Figure 9

1.40

APPARENT RESISTIVITY CAVE MODEL (THEORETICAL) DEPTH TO TOP = .A R„

DETECTIOil THRESHOLD FOR 20Z ERROR IN DATA

T.20

A 20% geologic noise level is probably the minimum to be expected due to changes in structure not related to the target cave. Thus if the size of the departure of the normalized apparent resistivity curve from 1.0 were much smaller than in Figure 10, the cave could not be detected. A deeper cave would, of course, give a smaller anomaly. Note that the maximum response occurs for an "a" spacing on the order of the depth (1.5 R ) to the center of the cave. Similar results for a spherical cave model are given by Habberjam (1969).

The electrical resistivity method has been shown to respond to zones of fracture concentrations revealed by fracture traces mapped on aerial photographs in carbonate areas in Centre County, Pennsylvania. Kirk (1976) found a decrease of apparent resistivity over Miller Cave using a 10-ft spacing. This "a" spacing is too small to result in measurements which are influenced by the cave which is 30 to 40 ft below the surface. The observed response was attributed to the fracture zone above the cave and filling of the cracked rock zone with loose soil, which would be more con­ductive than the limestone.

Johnson (1966) ran Wenner profiles over several fracture traces. Results are shown in Figures 11 and 12 which show that in some cases the zone of fracture concentration below the fracture trace will give an in­crease in the apparent resistivity curve, while in other cases a decrease occurs. The increases tended to occur for profiles run during dry weather and the decreases for profiles run after rain when soil water content is high. The interpretation made was that after a rain, soil above the frac­ture zone will be wet for a period of time and will be electrically conductive. After a period of time the soil above the zone will have been

3000

2000

1000

ELECTRICAL LES ! ST 1VITY PROFILE WEI1NER ARRAY, 50 FT A SPACRIO

DRY HEATHER

AFTER JOHNSON (1966)

FRACTLR TRACE

H h1 00 FT

LOCATIONS

1/ ^

DISTANCE — - »

Figure 11

476

2000

ci

GO

CO LiJ

1000 •

ELECTRICAL RESISTIVITY PROFILE WENNER ARRAY, 50 FT A SPACING

WET WEATHER AFTER JOHNSON (1966)

100 FT

IS FRACTURE TRACE LOCATION

DISTANCE

Figure 12

drained by seepage into the fracture system, leaving a dry electrically resistant soil. Unfortunately, Johnson (1966) did not repeat individual profiles several times to determine the period after a rain for which soils beneath fracture traces remain wet and conductive. His results do, however, point up the fact that the electrical resistivity method requires a good knowledge of the electrical characteristics of the area being monitored and how these characteristics change with rainfall history.

Electrical earth resistivity surveys, like geothermal surveys, are helpful in selecting which of many possible fracture traces mappable in a given area using air photographs delineate the most decomposed and permeable bedrock.

Seismic Methods An obvious seismic method for detection of underground cavities would

appear to be reflection seismology. Thus several attempts have been made to use seismic reflections to locate caves and tunnels. A discussion of this work is given by Dean (1975). None of the attempts have conclusively demonstrated that the method can be used to reliably detect a cavity. Cavities more than a few cavity diameters below the surface will not effi­ciently reflect low frequency seismic waves (wavelengths greater than the cavity diameter); high frequencies tend to be attenuated rapidly and thus the high frequency reflections will be small and lost in natural or source generated noise.

For cavities less than a cavity diameter below the surface, very high frequencies could be used. However, since the source and receiver must be close together, the source generated noise will probably make it impossible to see the reflection. Use of multiple geophone arrays may make it possible

477

SOURCE RECEIVER GEOMETRY

H—10 ft

o o

O - DETECTOR STAT I Of;

$ - SOURCE LOCATIO:;

After Hawk (1972)

Figure 13

to see reflections from very shallow cavities. However, there are many problems associated with combining high frequency signals from an array of geophones. These problems are magnified when the reflector is very near both the source and the receivers.

Another method for near surface cavity detection was attempted by Watkins et al(1967). In this work observations were made which may have been due to a resonance of the cavity. The period of the resonance sug­gested is set by the time a Rayleigh-type wave takes to go around the circumference of the cavity.

Here we will discuss an alternative technique for locating either a cavity or a fractured zone. This approach requires the use of a source at a point removed from the cavity, which generates Rayleigh-type surface waves. The waves then propagate along the surface to the cavity where the presence of the cavity disrupts them.

A series of field measurements were made by Hawk (1972) around typi­cal central Pennsylvania carbonate fracture zones. Figure 13 shows the source and a group of receivers in relation to a fracture trace. Two types of sources were employed; an impulsive source (a sledge hammer) and a harmonic wave source (a vibrating engine; seismic wave output was 25 Hz).

Figure 14 shows the ground amplitude versus receiver position for three profiles for the harmonic source. Notice the high amplitude that occurs over the fracture trace. This increase in amplitude over the trace was evident on the majority of profiles obtained by Hawk (1972). Thus a zone of high amplitude appears to be a good indicator of a fracture zone.

The transient source experiments also employed the source-receiver geometry of Figure 13. Two features of the seismic records (Figure 15) were indicative of the fracture zone. The first is the sudden disappear-

478

AMPLITUDE VS DISTANCE (HARMONIC, 25 Hz, SOURCE)

l o — T T T 1 -100 ï:'ÏS'i 70

PROFILE 7 (DISTANCE FROM FRACTURE IH FEET)

INDICATES

FRACTURE TRACE

PROFILE 5 (DISTANCE Z9

FACTURE III FEET)

Afte r Hawk (1972)

Figure 14

ance of the higher frequency energy for receiver positions at and beyond the fracture trace. The second feature is the greater time duration of the signal at receivers beyond the fracture trace compared to receivers between the source and the fracture trace.

The presence of the fracture zone may affect the Rayleigh wave in several ways. The soil in such a zone may have physical properties which are different from the soil over unfractured bedrock. Also, the bedrock will be fractured and have a lower strength than the unfractured bedrock. Travel-time studies (Lavin et al, 1972) have shown the seismic velocity to be lower in these zones. This lower velocity (thus lower elastic mod­uli) material can cause the increase in amplitude observed in the harmonic measurements.

The loss of the higher frequency waves in the transient experiments is due to one of two causes. First, the material within the fracture zone is less homogeneous than material adjacent to the fracture zone. This in-homogeneous material will cause loss of energy by scattering of energy out

479

of the Rayleigh wave. This scattering is more effective in attenuating high rather than low frequency waves. The second possibility is that the material differences only extend to shallow depths below the surface. The short period surface wave amplitudes die off more rapidly with depth than the longer period waves. If the wave decay depth is much larger than the depth to which the material is altered by the fracture, the wave will pass with little effect. Thus the short period waves will be strongly affected while the longer period waves will not.

We prefer the scattering hypothesis as the mechanism for attenuating the short period component of the waves. In addition to decreasing the high frequency waves, scattering will cause a lengthening of the signal at receivers beyond the fracture zone (Dainty et al, 1974).

This method of using surface waves for detecting fracture zones might be extended to the detection of cavities. To examine this possibility we performed a seismic model experiment. The model setup is shown in Figure 16. Cavities of two sizes were cut into a Plexiglas sheet. The cavities are shown in Figure 17. The difference between cavities D and E is that cavity E comes within 0.5 cm of the surface and cavity D stops 2.0 cm below the surface. The amplitude decay depth of the Rayleigh waves used in the model experiment (^15,000 Hz center frequency) was approximately

Figure 15

480

SOURCE V

2 DIMENSIONAL SEISMIC MODEL

RECEIVERS^ I—DISTANCE-^

80 cm

Figure 16

CAVITY — a

CAVITY SHAPES

A NO CAVITY

é 2.0cm

4.0 cm

~1~ 0.5 cm 0\

//,

Figure 17

5 cm. Thus the bottoms of cavities D and E were deep enough to keep most surface waves energy from passing below the cavity. Therefore, the energy had to reach the receivers beyond the cavity by going between the top of the cavity and the surface. Figure 18 shows the vertical component wave amplitudes for surface receivers. Curve A is for no cavity present. The decay with distance for curve A is just due to natural attenuation of the Plexiglas and should be considered as the standard for comparison with curves D and E. Notice that for receivers located between the source and the cavity the amplitudes for curves A, D and E are similar. However, the

481

< UJ 0 .

10

0

SEISMIC MODEL RESULTS AMPLITUDE VS DISTANCE

REFERENCED TO CENTER OF CAVITY

POSITION OF _ V. CAVITY CENTER 1 I

• i S » i -20 -15 -10 -5 0 +5 +10 +15 +20

DISTANCE [cm] AWAY FROM SOURCE

Figure 18

amplitude ratio of D to A is 0.6 and of E to A is 0.12 beyond the cavity, showing that the cavity will block the passage of the Rayleigh wave. Therefore, a sudden decrease in amplitude of the Rayleigh wave with dis­tance from the source will be indicative of a shallow cavity.

Conclusions Each of the four geophysical methods considered will, under favorable

conditions, indicate the presence of a void or a fracture zone in carbonate bedrock. The gravitational method is probably the most likely to give an anomaly over the caverns. However, the electrical and seismic methods may allow large areas to be surveyed more rapidly than the gravitation method and may thus have an economic advantage.

In all four geophysical methods the anomaly indicating the presence of the cave or fracture zone is due, in a major degree, to changes in the soil and in its moisture content above the feature.

To make proper use of any of the geophysical methods it is necessary to have an understanding of the normal geophysical response of the area in question and to be able to predict the form of the geophysical anomaly due to the void or fracture. This is particularly true when dealing with high "noise" areas.

Further, where zones of fracture concentration control weathering of carbonate rocks, these features should first be mapped using conventional aerial photography (when mapping fracture traces) and LANDSAT, SKYLAB imagery or U-2 level photography (when mapping lineaments. Geophysical surveys may be planned following this initial work to narrow down points of interest or potential problem areas. Geophysical surveys may then be fol­lowed up with test borings that can be located to provide maximum informa­tion with the least amount of drilling expense. Photo and imagery analysis

482

techniques are especially rapid and rather inexpensive, hence are a good starting point for planning most geophysical surveys in karst terrains. However, where modifications in land use have obscured the subtle evidence for fracture traces and aerial photographs that predate these changes do not exist, one will be forced back to a grid system application of geophys­ical surveying or a more random method of test drilling.

Acknowledgement Much of the research described in this paper was supported by the

Mineral Conservation Section, College of Earth and Mineral Sciences, The Pennsylvania State University. Computer results were done at the Computa­tion Center of The Pennsylvania State University. The seismic analog model experiments were carried out by R. Christen.

References

Arzi, A. A., 1975, Microgravimetry for Engineering Applications: Geophysi­cal Prospecting, v. 23, p. 408-425.

Bair, E. S., 1976, Permeability Distribution in Surficial Glacial Outwash Revealed by a Shallow Geothermal Prospecting Technique: M.S. Thesis, Dept. of Geosciences, Penn. State Univ., University Park, PA, 91 pp.

Bauer, L. D., W. H. Gardner, and W. R. Gardner, 1972, Soil Physics: John Wiley and Sons, Inc., New York, NY.

Cartwright, K., 1968, Thermal Prospecting for Groundwater: Water Resources Res., v. 4, p. 395-401.

Colley, G. C , 1963, The Detection of Caves by Gravity Measurements: Geo­physical Prospecting, v. 11, p. 1-9.

Dainty, A. M. , M. N. Toksb'z, K. R. Anderson, P. J. Tines, Y. Nakamura, and G. Latham, 1974, Seismic Scattering and Shallow Structure of the Moon in Oceanus Procellarum: The Moon, v. 9, p. 11-29.

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