DEPARTMENT OF CONSERVATION-
                Michael F. Byrne, Director
                                           PB95-191268
                                          . EPA No. 530-R-95-013a
              Application  of Geophysics
                                 to
       Acid Mine Drainage  Investigations
                             Volume!
                        -
                        literature Review
                         '"- >< , -' „ «n/f
                        ,  -
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                             Table of Contents
                                  Volume I
        ^

                                                                      Page

      EXECUTIVE SUMMARY                                             v.

      RECOMMENDATIONS                                             vii.

I.     INTRODUCTION
            PURPOSE AND SCOPE                                        1.
            ACKNOWLEDGEMENT                                        2.

II.    ACID MINE DRAINAGE (AMD)
            GENERAL DISCUSSION OF AMD                               5.
            GEOCHEMICAL SETTING FOR AMD                            5.
            GEOCHEMICAL AND GEOPHYSICAL RELATIONSHIPS
                        General Discussion                                  6.
                        Porosity, Permeability and Heterogeneity       •           11.
                        Electrolyte Concentration                             14.
                        Moisture Content                                   19.
                        Temperature                                      21.

III.    APPLICATION OF GEOPHYSICAL METHODS TO AMD INVESTIGATIONS
            GENERAL DISCUSSION                                      23.
            OPTIMUM SAMPLING IN GEOPHYSICAL SURVEYS              24.
            D.C. RESISTIVITY GEOPHYSICAL METHODS
                  GENERAL DISCUSSION
                        Basic Principles                                    26.
                        Equivalence and Suppression                          27.
                        Survey Methods                                    29.
                        Array Electrode Geometries                           30.
                        Depth of Investigations                              32.
                        Comparison of Arrays                               35.
                        Interpretation of Resistivity Data                        36.
                  USE OF D.C. RESISTIVITY IN AMD INVESTIGATIONS
                        General Discussion                                 39.
                        Geologic Structure and Stratigraphy                     40.
                        Ground- Water Protection Zones                        41.
                        Aquifer Properties and Ground-Water Quality              41.
                        Fracture Systems                                   42.
                        Landfills and Waste Impoundments       -              44.
                        Open and Filled Voids                              44.
                        Long-Term Monitoring                              45.
                  D.C. RESISTIVITY FIELD PROCEDURES                 45.
                  MISE-A-LA-MASSE GEOPHYSICAL METHODS             47.

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                         Table of Contents
                            Volume I
                                                            Page
          INDUCED POLARIZATION (IP) GEOPHYSICAL METHODS
               GENERAL DISCUSSION                            48.
               USE OF IP IN AMD INVESTIGATIONS                  52.
          ELECTROMAGNETIC (EM) GEOPHYSICAL METHODS
               GENERAL DISCUSSION                            54.
               VERY LOW FREQUENCY METHODS
                    Conventional VLF                             57.
                    VLF Resistivity                                58.
               AUDIO-MAGNETOTELLURIC AND CONTROLLED
                    SOURCE AUDIO-MAGNETOTELLURIC METHODS   59.
               FREQUENCY-DOMAIN SLINGRAM
                     AND GROUND CONDUCTIVITY METER METHODS 60.
               TIME-DOMAIN EM METHODS                       64.
               USE OF EM METHODS IN AMD INVESTIGATIONS        65.
               EM FIELD PROCEDURES                           70.
          SELF POTENTIAL (SP) GEOPHYSICAL METHODS
               GENERAL DISCUSSION                            71.
               USE OF SP IN AMD INVESTIGATIONS                 74.
               SP FIELD PROCEDURES                            76.
          SEISMIC GEOPHYSICAL METHODS                        77.
          GRAVITY GEOPHYSICAL METHODS                       84.
          MAGNETIC GEOPHYSICAL METHODS                      87.
          GROUND PENETRATING RADAR METHOD                  88.
          BOREHOLE GEOPHYSICAL METHODS                      89.

IV.   REFERENCES
          GEOPHYSICAL TEXTBOOKS                             95.
          GENERAL ENVIRONMENTAL, GROUND WATER
               AND GEOTECHNICAL GEOPHYSICAL METHODS        97.
          WATER QUALITY AND GEOPHYSICS                      100.
          CONVENTIONAL RESISTIVITY GEOPHYSICS                 104.
          MISE-A-LA-MASSE GEOPHYSICS                          118.
          INDUCED POLARIZATION GEOPHYSICS                    119.
          ELECTROMAGNETIC GEOPHYSICS                        122.
          SELF POTENTIAL GEOPHYSICS                           127.
          SEISMIC GEOPHYSICS                                  130.
          GRAVITY GEOPHYSICS                                 133.
          MAGNETIC GEOPHYSICS                                135.
          GROUND PENETRATING RADAR                         136.
          BOREHOLE GEOPHYSICS                               138.

V.   LIST OF ABBREVIATIONS                                    143.
u.

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                             List of Figures
                                Volume I
                                                                            Page

1.      Plot of specific conductance versus total dissolved solids                   16.

2.      NaCl equivalence concentrations chart                                   16.

3a.    Potential and current distribution for resistivity array                      31.

3b.    Wenner array equipotential and current flow lines                        31.

4.      Electrode configuration D.C. resistivity with interpretation formulas         33.

5a.    Electromagnetic response versus depth curves                             63.

5b.    Electromagnetic cumulative response versus depth curves    .              63.

6.      Schematic of seismic refraction ray-paths                                 79.

7.      Field setup of shot points and geophones for seismic refraction             79.

8.      Seismic ray-paths and time-distance plot for 2-layer model                 82.

9.      Matrix chart of borehole logging methods applicability                     94.



                                 Tables
                                Volume I


                                                                            Page

1.      Typical AMD contaminant levels                                         7.

2.      Resistivity-conductivity conversion factors                                  9.

3.      Resistivities of selected rocks and minerals                                10.

4.      NaCl conversion coefficients                                            15.

5.      Ionic conductivity and mobility                                          18.

6.      Summary of EM survey methods                                        55.
                                                                                 111.

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                              EXECUTIVE SUMMARY
          ,i
Surface and borehole geophysical methods are an important tool in the investigation and
monitoring of acid mine drainage (AMD) pollution.  Geophysical methods can be used in AMD
investigations to:

      o   map sites of disposal of mining waste;
      o   map the extent of AMD pollution in aquifers;
      o   monitor velocity of AMD ground-water pollution;
      o   delineate depth and lateral extent of waste fills;
      o   identify potential source areas of AMD;
      o   locate geologic structures such as faults and formation contacts;
      o   identify gravel channels within finer grained materials;
      o   define water table in unconsolidated sediments;
      o   identify subsurface voids;
      o   map higher permeability areas within aquifers;
      o   define the extent of weathering, fracturing and faulting;
      o   define the extent of protective clay layers;
      o   map soil conductivity; and
      o   provide complimentary data for correlating borehole and monitoring well data.

Geophysical methods that can directly measure AMD pollution rely on the resistivity or
conductivity contrast caused by the increase in total dissolved solids associated with the lowering
of pH and the dissolution of sulfide minerals. This increase in soil/rock conductivity is due to a
large increase in specific conductance of ground water within the soil/rock pores.  The
relationship between the ground-water geochemistry and the resistivity of the whole soil/rock
mass is complex and non-linear.  State-of-the-art application of geophysical methods to
investigations of environmental and hydrogeologic problems requires that site-specific data be
developed to calibrate  general empirical relationships. Geophysical techniques best suited to
directly detect subsurface AMD pollution include direct current (D.C.) resistivity and
electromagnetic methods.  Induced polarization, self potential, mise-a-la-masse methods may also
be of value in some hydrogeologic settings. Although other geophysical methods do not directly
measure conductivity, they can be of value in AMD investigations in that they provide
complementary information to aid in the subsurface interpretation. These methods include:

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seismic, gravity, magnetometer, ground penetrating radar, and various borehole methods.


Field investigations to evaluate the application of D.C. resistivity, electromagnetic, self potential

and magnetometer methods were undertaken at four closed mine sites with known or suspected

AMD pollution. Results of the field investigations are presented in the second volume of this

report. The following conclusions were drawn from an analysis of the survey data collected in

the field investigations.


o    Electromagnetic (EM) surface methods were successful in detecting and mapping acidic
     ground water in mine waste piles.

o    D.C. resistivity methods were successful in developing vertical profiles of acidic mine waste
     material that correlated well with electromagnetic survey data and well logs.

o    Self potential surveys were successful in detecting acidic ground water flow from mine waste
     ponds.

o    Magnetometer surveys were proved useful in distinguishing between an increase in
     subsurface conductivity due to buried man-made iron or steel objects, and higher specific
     conductance from acidic ground water or high conductivity of soils,  rock or mine waste.


Included in this report is a review of the literature related to the reported and potential

application of geophysical methods to AMD investigations. The report also briefly discusses:


o    The geochemistry of AMD.

o    The relationship between ion concentration and specific conductance.

o    Empirical relationships that are available to predict the resistivity of soil/rock.

o    Formulas for determining the optimum line spacing for geophysical surveys and the
     associated probabilities.
 VI.

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                               RECOMMENDATIONS
The following are recommended after an extensive review of literature on geophysical
applications to acid mine drainage (AMD) and mapping of fluids with high specific conductance:

o    A combination of several geophysical methods such as EM, resistivity, self-potential and
     seismic provide a cost-effective method for evaluating potential impacts, defining subsurface
     structures, identifying potential flow paths for contaminants, and defining the extent of
     ground-water pollution.

o    Geophysical surveys should become a standard procedure for preliminary investigation and
     monitoring of AMD where subsurface pathways are not readily discernable.

o    Geophysical surveys should be used to  monitor the migration of AMD plumes between
     points of sample collection, such as monitoring wells and springs, to confirm that the best
     sample locations have been selected  and to help in understanding seasonal changes in
     ground-water movement.

o    Geophysical instruments are expensive and require technical training to use. Many local
     and state agencies faced with AMD do not have the economic resources and staff to
     develop an in-house geophysical program. Ready access to geophysical expertise is
     desirable.  Consideration should be given to developing regional groups of geophysical
     experts that can provide professional services and  training to local and state agencies.
     These regional groups should reside in organizations that have a strong interest in applied
     geophysical research. Such organizations might include universities, the U.S. Geological
     Survey, some state geological surveys, or technical committees composed of geophysical and
     geologic professionals from government and the private sector..
                                                                                      vu.

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                                I. INTRODUCTION
                            PURPOSE AND SCOPE OF STUDY

Acid mine drainage (AMD) contributes to the pollution of both surface and ground water.
Increasing demand on limited water resources necessitates identification and monitoring of
sources of contamination such as AMD. Traditional methods used for detection, monitoring and
delineation of ground-water contamination rely on monitoring wells.  Proper location of
monitoring wells is critical for obtaining accurate data from which to interpret the extent and
concentration of contaminants. A misplaced well can yield erroneous data on the impact of the
contaminants and increase long-term monitoring costs.

AMD is generally associated with an increase in concentration of heavy metals and other ionic
species which increase the specific conductance of both surface and ground waters.  This increase
in specific conductance allows mapping of AMD ground-water contamination using geophysical
methods such as resistivity, electromagnetics, and self potential. The conductivity of mine waste
and underlying material is a function of the type of soil and rock, the porosity, and the specific
conductance of the fluids that fill the pores.  The specific conductance of the pore fluids is often
the dominant source of the electromagnetic response (McNeill, 1980). Electrical geophysical
methods can be used to map inorganic contaminants, identify direction and extent of contaminant
flow, estimate concentration gradients, develop time-series measurements of plumes, and provide
data for contaminant plume modeling.  Although geophysical studies should not completely
replace  the use  of  monitoring wells to collect subsurface hydraulic and water-quality data, they
can provide a low cost alternative to using only monitoring wells for the detection of AMD and
help in selecting locations for monitoring wells.

For this report several tasks were undertaken to evaluate the utility of geophysical techniques in
detection and monitoring of AMD contamination from mine wastes.  The tasks were conducted
in three phases:

o    Review  and summary of literature on geophysical methods that may be useful  in evaluating
     migration  of the high specific conductance contaminants in  ground water.
                                                                                       1.

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o    Selection of several closed mine waste sites in California as candidates for site-specific
     geophysical study.
          x
o    Evaluation of selected sites using electrical resistivity, self potential, and electromagnetic
     geophysical surveys to determine the applicability of these techniques in detecting and
     evaluating AMD contamination from mine waste.    Sites with existing ground-water-quality
     data were of particular interest because the data allow a semi-quantitative correlation
     between geophysical response and ground water quality.

The geophysical and geochemical criteria identified in the literature review, in conjunction with
certain non-technical criteria such as site  access, were used to develop system to rank mine sites
for site-specific investigation.

Reviewed literature focused on geophysical and geochemical conditions that are known to be
associated with AMD and can be detected by surface and borehole geophysical techniques.
These conditions include:

o     The presence of conductive minerals, such as pyrite, chalcopyrite, pyrrhotite, hematite, etc.
o     The presence of conductive dissolved electrolytes in surface and ground water, such as H+,
      Fe2+, Fe3*, SO^, Cu2+, and Zn2+.
 o    Porosity of soils and interconnection of rock fractures.
 o    The extent to which pores are filled with water, i.e., moisture content.


                                 ACKNOWLEDGEMENTS

 This report was prepared under a grant from the U.S. Environmental Protection Agency through
 the Western Governors' Association (Grant X-820497-01-0). The project was administered by
 Mr. Stephen Hoffman of the Office of Solid Waste and Emergency Response, U.S.
 Environmental Protection Agency. The project was conducted under the guidance of Messrs.
 James S. Pompy and Dennis O'Bryant of the California Department of Conservation, Office of
 Mine Reclamation (DOC-OMR).  Technical review and assistance was provided by Messrs.
 Michael Hunerlach, Charles Alpers and F. Peter Haeni of the U.S. Geological Survey, and Dr.
 Rodger Chapman of the California Division of Mines and Geology (retired). Field assistance
 was provided by Ms. Catherine Gaggini and Mr. Steve Newton-Reed of the DOC-OMR, and
 Messrs. Michael Hunerlach, William Hardy and Scott Hamlin of the U.S. Geological Survey,
 2.

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Water Resources Division, California District in Sacramento.  Special thanks go to Mr. William
Croyle of the Central Valley Regional Water Quality Control Board, Ms. Catherine Schoen of
the Tahoe Regional Water Quality Control Board, Mr. Rick Sugarek of Region IX USEPA, and
Mr. Steve Muir of WZI Corporation for their assistance, without which this project would not
have been possible.

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                      II.  ACID MINE DRAINAGE (AMD)
                           GENERAL DISCUSSION OF AMD1

When iron sulfide minerals are exposed to air and water by mining activities, they oxidize and
can produce sulfuric acid. The low pH water in turn leaches metals from other minerals
associated with the sulfide mineralization. The problem of AMD from mining is complicated by
time delays associated with the formation of AMD.  Time delays in AMD production are often
due to eventual consumption of gangue minerals with buffering capacity, usually calcite and other
carbonates, as well as to changes in the nature of ore with depth (e.g. increasing sulfide content).
Often, mines with no problem during early periods of mining develop AMD several years or
decades later.

Conclusions regarding AMD reached in a Mine Waste Study commissioned by the State Water
Resources Control Board and conducted by the University of California at Berkeley are
(CSWRCB,  1988):

o    AMD is produced by naturally occurring minerals that often are present in great quantities
     in mine wastes.
o    Once AMD develops, it tends to get worse. Common iron and sulfur oxidizing bacteria,
     Thiobacdllus and  Ferrobacillus femxxddans, accelerate the process when the pH of the mine
     waters falls below a critical value. Thus, the process can be difficult to stop.
o    AMD may not develop during mining operations but after closure, when mitigation
     measures are difficult to adopt.
                          GEOCHEMICAL SETTING FOR AMD

AMD is generated from rock containing sulfide minerals, in particular iron sulfides such as pyrite
and marcasite, that are exposed to oxygen by air and water during mining. Stumm and Morgan
(1981) list the following reactions as characteristic of the oxidation of pyrite when exposed to air
and water:
        References begin mi page 100.

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                    FeS2(s) + 7/2 02 4- H2O = = > Fe2+ + SO42' + 2 H+              (1)
                    Fe2+ + 1/4 O2 + H+ = = > Fe3+ + 1/2 H2O                       (2)
                    Fe3+ + 3 H2O = = > Fe(OH)3(s) + 3 H+                        (3)
                    FeS2(s) + 14 Fe3+ + 8 H2O = = > 15 Fe2+ + 2 SO42- + 16 H+     (4)

AMD generation is accelerated by acidophilic bacteria, Thiobacillus ferrooxidans and Thiobacillus
thiooxidans, which derive energy from the oxidation of sulfide minerals and Fe2+, accelerating
reaction (2) by  3 to 6 orders of magnitude (CSWRCB, 1988).  The production of acid is greatest
at about pH 3.0 due to bacterial activity and diminishes between pH 3.0 and 2.0 due to buffering
reactions that precipitate sulfate, ferric sulfate, and ferric hydroxides (Singer and Stumm, 1970;
CSWRCB, 1988).

The oxidation of sulfides and resulting generation of acid releases heavy metals  and sulfosalts
from ore-forming minerals and gangue rock. Heavy metals such as silver (Ag), cadmium (Cd),
cobalt (Co), copper (Cu), mercury (Hg), manganese (Mn), molybdenum (Mo), nickel (Ni), lead
(Pb), zinc (Zn), arsenic, (As), antimony (Sb), and selenium (Se) are commonly associated with
AMD.  Table 1 lists typical ranges of AMD contaminant levels from the Appalachian coal region
of the United States (Onysko, 1985). Data compiled as part of a study of mine waste in
California (CSWRCB, 1988) found the concentration of heavy metals rises with a decrease in pH
and, except for copper, zinc, and sulfate, concentration of most other heavy metals and sulfosalt
anions rarely exceeds 100 ppm (CSWRCB, 1988).
                 GEOCHEMICAL AND GEOPHYSICAL RELATIONSHIPS

General Discussion

The ability of geophysical methods to detect and map an AMD plume in the subsurface is
directly related to the contrast between background and AMD plume conductivity. The degree
of contrast is controlled by the signal noise caused by variability of the hydrogeology, cultural
features, and by limitations of the instruments in sensing plumes at depth. Benson and others
(1985) evaluated the correlation between geophysical signature of high specific conductance
ground-water plumes and the quality of ground water.  They evaluated sources of high specific
conductance plumes caused by  chloride, sodium, and sulfate that leaked into shallow ground

6.

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                                 Table 1
         TYPICAL ACID MINE DRAINAGE CONTAMINANT LEVELS2
Parameter
    pH
Iron
    Fe2+
    Fe3+
    Total Soluble, as Fe
Sulfate
    so,2-
    HSCv
    Total, as SO42-
Aluminum
Manganese
Acidity to pH 3, as CaCO3
    Hydrogen Ion
    Bisulfate Ion
    Ferric Ion
    Aluminum Ion
    Ferrous Ion
    Humic Acids
    Total
Total Dissolved Solids
Suspended Solids
Ionic Strength, molar
Redox Potential
Range of Concentration
8.3
0
0
0
50
0
50
0
0
0
0
0
0
0
0
0
250
5 .
0
-200
- 1.3
- 5,000 mg/L
- 1,000 mg/L
- 6,000 mg/L
- 10,000 mg/L
- 20,000 mg/L
- 20,000 mg/L
- 100 . mg/L
- 100 mg/L
- 2,500 mg/L
- 10,000 mg/L
- 1,000 mg/L
- 500 mg/L
- 5,000 mg/L
- 500 mg/L
- 20,000 mg/1
- 25,000 mg/L
- 5,000 mg/L
- 0.7
- +1,000 mV
' Table after Onyska, 1985

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water.  D.C. resistivity and electromagnetic (EM34-3 and EM31) methods were compared.
Specific conductance measured in ground-water monitoring wells within the plume and in
background areas were found to have a correlation coefficient ranging from 0.77 to 0.90 at a 95
percent confidence interval with the surface geophysical measurements.  Increases in
hydrogeologic complexity, the number of cultural features, and depth of plume all contributed to
a reduction in correlation. Researchers have found that a plume conductivity of at least twice
background is necessary to identify a plume boundary (Greenhouse and Slaine, 1986; Grady and
Haeni, 1984;  and Slaine and Greenhouse,  1982 ).

Correlation of electrical geophysical instrument readings and ground water quality can be done
either by using the bulk material conductivity with empirical equations, such as Archie's law, or
by site specific comparison of geophysical readings with water quality data (Benson and others,
1985; and Grady and Haeni, 1984 ).  The latter provides a more reliable correlation while the
former can help predict variations in readings.

Electric current can flow through a geologic material in three ways: ohmic, electrolytic, and
dielectric conduction (Telford and others,  1990). Ohmic current occurs in materials such as
metals where free electrons can easily be made to flow. Electrolytic current occurs with ions in
solution and is comparatively slow.  Dielectric conduction occurs in poor conductors or insulators
when an external electric field causes electrons to displace slightly, resulting in a temporary
polarization of the material.

Electrical resistivity is a measure of the ease with which electrical current flows through a
substance. Resistivity (p) of a material of cross-sectional area (A), length (L) and resistance (R)
between two faces is:
                                       p = R-A/L

where:                            R  =  V/I (Ohm's law)
                                   V  =  voltage, volts
                                   I  =  current, amperes

Electrical resistivity is a property of the medium only and is not proportional to  the dimensions
of the medium.  The units of resistivity in the MKS system (meter-kilogram-second) are ohm-
meters and in the COS system (centimeter-gram-second) ohm-centimeters. Conductivity (a) is

8.

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the reciprocal of resistivity (a =  1/p), and is commonly measured in milliSiemens per meter
(mS/m) or rnillimhos per meter (mmhos/m), where 1 Siemen *= 1 mho.  Historically, the
conductivity of water, or specific conductance, is given in units of millimhos per meter (mmho/m),
micromhos per centimeter (/imhos/cm). More recently unit of Siemen has replaced the mho.
Table 2 is a list of conversion factors for resistivity to conductivity.
                                         Table 2

                 RESISirVTIY-CONDUCTTvTTY CONVERSION FACTORS

                           1 ohm-meter =  100 ohm-centimeter
                             1 / ohm-meter = 1 mho / meter
                                     1 / ohm  =  mho
                                   1 mho =  1 Siemen
                        1 mho / meter =  1000 milliSiemens / meter
                  1 mho / centimeter   =  1 x 106 microSiemens / centimeter
                     1 / ohm-meter =  1 x 104 microSiemens / centimeter
                   1 milliSiemen / meter =  10 microSiemens / centimeter
Conductivity of most soil and rock minerals is generally very low because most are electric
insulators.  The presence of some minerals such as magnetite, specular hematite, carbon,
graphite, pyrite and pyrrhotite can greatly increase the conductivity of rock and soils (Keller and
Frischknecht, 1966).  Table 3 lists resistivity values for selected soil and rock minerals.

Conductivity of electric current through soil and rock is primarily through the pore water
electrolytes. Thus, the volume of pores, the interconnection of pore space and the characteristic
of the pore water electrolyte are all important factors in determining the overall conductivity of
soil or rock. The following are factors that affect conductivity or resistivity of soils and rocks and
are discussed below.

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                                        Table 3

                RESISTIVITIES OF SELECTED ROCKS A>4D MINERALS3
     .  Rock/mineral-
Low
Resistivity (Q • m)
     High
Average
Arsenopyrite, FeAsS
Qornite, CujFeS4
Chalcocite, Cu,S
Chalcopyrite, CuFeSj
Copper-native, Cu
Coveliite, CuS
Cuprite, Cu,O
Enargitc, GijAsS,
Galena, PbS
Hematite, FejOj
Ilroenite, FeTiO,
Iron-metal, Fe
Magnetite, Fe,O4
Marcasite, FeSj
Pyrite,FeSj
Pyrrhotite, Fe^S,
Sphalerite, ZnS
Sulfur, pure crystals, S
Tetrahedrite, Cu10(Fe,Zn)7Sb4S13
Argillites
Basalt
Diabase
Gabbro
Gneiss
Granite
Limestone
Marble
Quaternary/Tertiary terrestrial sands
Quaternary/Tertiary marine sands
2.0 x 10 5
1.6 x 10*
8.0 x ID'5
3.0 x 10s
1.2 x 10*
3.0 x 10'
1.0 x 10*'
2.0 x 10*
6.8 x 10*
2.1 x 103
1.0 x 10 J

1.5 x 10s
1.0 x Ifr3
6.0 x 10s
2.0 x 10*
2.7 x ID"3

3.0 x 10 '
1.0 x 10*'
1.0 x 10*'
2.0 x 10*'
1.0 x 10*3
6.8 x 10* VO
3.0 x 10*2
5.0 x 10*'
1.0 x 10*2
1.5 x 10*'
1.0 x 10°
1.5 x 10*.'
6.0 x 10 3
1.0 x 10""
83 x 10*'
3.0 x 107
8.0 x 107
5.0 x 10+1
9.0 x 10 '
5.8 x 10 '
4.0 x 10°
4.0 x 10°

1.0 x 10"
1.5 x 10-'
1.2 x 10°
1.6 x 10^
4.0 x 10+<

3.0 x 10*4
SxlO1
13 x I0+7(diy)
5.0 x 10*7
1.0 x 10*'
3.0 x 10+6(w«)
1.0 x 10*'
1.0 x 10+7
2.5 x 10*'(diy)
• 5.0x10*'
1.0 x 10*'
1.9 x 103
9J.x 10s

7.3 x 10^



2.0 x 10 2
1.9 x 102


1.0 x 10*
5.0 x 10*

2.2 X 10 '
7.7 x 1C"5

1.0 x 10*"
1.1 x 10*2










 Table after Keller, 1966 and 1989; and Telford and others 1976 and 1990
10.

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                                   Table 3 (continued)
                 RESISTIVITIES OF SELECTED ROCKS AND MINERALS
        Rock/mineral
                Low
                    Resistivity (Q • m)
                          High
Average
Quartizite
Quartz diorite
Sandstone
Shale
Shists, calcareous & mica
Slcarn
Slates
Tuff
Unconsolidated alluvium & sands
Unconsolidated clays
1.0 x 10*'
2.0 x 10*<
1.0 x 100
2.0 x 10+I
2.0 x 10+1
2.5 x 10+2(wet)
6.0 x 10**
2.0 x 10* V«)
1.0 x 10* '
1.0 x 10°
2.0 x 10*'
2.0 x 10* «(wet)
6.4 x 10+8
2.0 x 10°
1.0 x 10*'
2.5 x 10+'(diy)
4.0 X 10*'
1.0 x 10+Vy)
8.0 x 10*2
1.0 x 10*2









2.0 X 10M(»e<)
o    Porosity, permeability and heterogeneity
o    Electrolyte concentration
o    Moisture content
o    Temperature
Porosity, Permeability and Heterogeneity

Conductivity of clean sands can be expressed by the empirical formula of Archie (Telford and
others, 1990):

                                   Pe = a • 
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                     <|>      =      porosity, expressed as a fraction of pore volume
                     S      =      fraction of pores containing water
                     a,n,m  =      constants of regression, with the following ranges,
                                       0.5Sa<2.5
                                       1.3 < m < 2.5
Mazac and others (1987) note that "m", the coefficient of cementation, is somewhat larger than 2
for cemented and well sorted granular rock and somewhat less than 2 for poorly sorted and
poorly cemented granular rock. For unconsolidated, weakly cemented materials with a porosity
that ranges from 20 to 70 percent, a value of 1.3 can be assumed for "m" (McNeill, 1980; and
Keller, 1989).  The "a" coefficient varies from slightly less than 1 in rock with intergranular
porosity to slightly more than 1 in rock with only joint porosity and as high as 3.5 in vesicular
tuffaceous rock (Carmichael,  1989).

A commonly used term in formation evaluation is the "formation factor" (FF) which is equal to
Pe/Pw  (Telford and others, 1990).  Lynch (1962) provided an alternate formula for formation
factor stating that it is a constant of proportionality between the bulk resistivity  of the formation
and the resistivity of the pore water.  It depends only on the tortuosity and porosity of the rock
and is independent of the water in the pore space as expressed in the following formula:

                                       pe  =  FF-p

where:                             FF     =     (L./L)2 • (1/<|>)
                                   (IVL)2 =     tortuosity
                                   L      =     actual length of a rock block
                                   Le      =     length of tortuous path through pores in
                                                block, L < L,.
                                   <(»      =     porosity

For soils, the minerals of sand and silt grains are generally electrically neutral and act as
insulators.  Clays  act either as insulators or conductors depending on their moisture content and
cation exchange capacity (CEC). The CEC is a measure of the number of cations required to
neutralize the clay particles and is given in milliequivalents adsorbed per 100 grams of soil.

12.

-------
Clays develop from sheet-like or layered silicates which have a negatively charged surface.
Cations such as Ca2+, Mg2"1", H+, K+, Na+, NH4+ are loosely held to clay particle surfaces and are
available for exchange with other more strongly attracted cations, or they may go into solution
when water is present.  The small particle size of clays, < 0.002 millimeters in diameter, provides
a very high surface area per unit volume of soil. Thus a large number of cations can be adsorbed
per unit volume and it is these adsorbed cations that affect the soil conductivity.

Although clay particles are poor conductors, a small amount of moisture will create a thin layer
on the particle where adsorbed ions can be dissociated and become available for ionic
conductivity. For shaly sandstones, Lynch (1962) shows that an apparent formation factor FFa
can be estimated using a method proposed by Hill and Milburn (1956). Specifically, FF? can be
related to a formation factor for a fluid with a resistivity of 0.01 ohm-meters (FF0 01) by the
following formula:
                              FFa = FojM-OOO-p.)1"1*000 ^

where:                      b    ^=     - 0.135 •  (CEC/<|>) - 0.0055
                            <|>      =     volumetric water content, cm3 water/ cm3 soil
                            pw     =     resistivity of pore water

Although this formula was derived for formations encountered in oil exploration, it indicates that
CEC has an exponential effect on the conductivity of rock.  For soils and rocks where the pore
water is not saline, the conductivity can be strongly influenced by clay content and should always
be considered when interpreting geophysical surveys (McNeill, 1980).

McNeill (1980) proposed the following rules of thumb for determining the change in pore water
concentration needed to affect bulk soil conductivity:

                                      aa = 0.25 • aw
                                     aw = (1/6)  • TDS
                                     ca  = (1/25) • TDS

where:                      aa     =  apparent bulk soil conductivity, mS/m
                            ow     =  pore water specific conductance, mS/m
                            TDS    =  total dissolved solids, ppm

                                                                                       13.

-------
Field application of geoelectric methods requires consideration of the way electrical properties
are averaged for sections of soil and rock which are generally heterogeneous.  Only borehole
geophysical logging provides a method for measuring resistivity of individual layers. Surface
geophysical methods average weathering, depositional layering, and structural  discontinuities.
The geologic heterogeneity may also make the earth material electrically anisotropic,  that is, the
resistivity may vary with orientation. For example, stratified rocks are generally more conductive
along bedding planes than perpendicular to them.  Keller and Frischknecht (1966) point out that
a distinction should be made between the "geoelectric section" and the "geologic section".  Their
boundaries may differ because the geoelectric boundary is based on resistivity  contrasts which
may not follow a boundary defined on geologic principles such as the fossil record.
Electrolyte Concentration

For electrical current to flow through soil or rock, ions must move through the electrolyte
solution in the pores.  The specific conductance of an electrolyte is proportional to both the total
number of ions in solution and their velocity or ionic mobility. The term specific conductance is
used to express the electrical conductivity of a body of unit length and unit cross section at a
specific temperature.  Pure water has a very low specific conductance, a few hundred micromhos
per centimeter (/^mhos/cm) or microSiemens/cm (pS/cm) at 25° C (Hem, 1989). For natural
water an approximately linear relationship can  be found between the total dissolved solids and
the specific conductance.

An estimate of the relationship between specific conductance and total dissolved solids can be
expressed by the formula:
                                       S = K-A

where:                      K     =     specific conductance,/unhos/cm
                            S      ~     total dissolved solids, mg/L
                            A     -     regression coefficient

Figure 1, taken from Hem  (Fig. 10, p. 67,1989), shows a straight line with a slope of 0.59 visually
fitted to data for waters of the Gila River.  A linear regression fit to the lower concentration data
shows that the slope of the line is steeper.  A  regression line slope ranging from 0.54 and 0.96

14.

-------
can be expected for natural waters with most slopes lying between 0.55 and 0.75. The higher
values are associated with waters that have a high concentration of sulfate.

An alternative method for estimating specific conductance and quality of water is to use the
Schlumberger NaCl curves in Figure 2 developed for fluid-conductivity logs (Alger, 1966; Keys
and MacCary, 1971; and Keys, 1989).  To use this method, convert the ionic concentration, in
milligrams per liter, to an equivalent NaCl concentration, then sum to find a total equivalent
NaCl solution.  On Figure 2, use the diagonal line equal to the NaCl concentration to find the
temperature of the formation waters and read the conductivity or resistivity of the solution on the
top or bottom horizontal scales, respectively. The ion conversion factors for calculating an
equivalent NaCl solution are given in Table 4 (Keys, 1989; and Lynch, 1962).

For an electrolyte  to conduct electric current the solute ions must move through the solvent
(water) to transfer charge.  The effectiveness of an ion in transferring charge depends upon its
charge, its size, and the way it interacts with the solvent (Hem, 1989). A measure of the
potential for an ion to transfer charge can be found in its ionic mobility. The ionic mobility
represents the terminal velocity, in centimeters per second, of an ion in a potential gradient of 1
volt per centimeter (Barrow, 1979).   Ionic mobility is a function of temperature and
concentration of salts in solution. In a highly-concentrated solution an ion's velocity or ionic
mobility is reduced by the motion of other ions. Changes in temperature also affect the viscosity
of the solution and the ability of ions to move.  The effect of temperature changes on ion
mobility is discussed in more detail below. Table 5 lists the ionic mobility for  selected ions at low
concentrations and standard temperature of 25° C (Vanysek, 1993).

                                            Table 4
                               NaCl CONVERSION COEFFICIENTS4

                            Ion                                Conversion
                                                               Coefficient
                            Ca2+                                 0.95
                            Mg2*                            '    2.00
                            K+                                  1.00
                            SO4-                                 0.50
                            HCO3-                               0.27
                            C032'               •                 1.26
          4 Conversion coefficients from Keys and MacCary, 1971

                                                                                        15.

-------
    500   1000   1500   2000   25OO   3000


       DISSOLVED SOLIDS. IN MILLIGRAMS PER LITER
                                                                        3500
         Figure 1.    Dissolved solids and specific conductance of composites of daily samples

                      from Gila River, Arizona (from Hem, 1989).
         in
         «J

         in
         lit
         u
         oc

         e

         3
            60
                                        RESISTIVITY — ohm.m«t«rs


                                            5      10       20
                                                                    50
                                                                                   200
                                                                                         140
                 20000   10000
5000
                                           2000    1000     500


                                        CONDUCTIVITY — /jmhoi/cm
                                           100
                                                   50
          Figure 2.    Electrically equivalent concentrations of a sodium chloride solution as a

                       function of resistivity or conductivity and temperature (from Keys and

                       MacCary, 1971).
16.

-------
An approximation of the specific conductance for dilute solutions of natural water can be
calculated with the following formula (McNeill, 1980; Barrow, '1979; and Keller and
Frischknecht,1966):
                   a  = y- [ClMl + QMj + ...+ CnMn]  = 96500 • Z(QMj)
where:
a
Q
                                   specific conductance in mS/m
                                   gram-equivalent weight of the itb ion per cubic meter of
                                   water (gm-eq-wt-nr3). A gram-equivalent weight is the
                                   formula or molecular weight of an ion divided by its
                                   valence.
                                   mobility of the ith ion in square meters per second per volt
                                   (m'-sec-'-V1).
                                   Faraday's constant, 96,500 coulombs of charge per gram
                                   equivalent weight of salt in solution.  1  coulomb =
                                   ampere-sec
For example, one gram of common salt, NaCl, dissolved in 1 cubic meter of pure water (1 mg/L
concentration) would have a specific conductance of 0.22 mS/m = 2.2 /uS/cm (McNeill, 1980).

               a  =  96500 • {[(23/58)/23] • 5.2 x Hr8 + [(35/58)/35] • 7.9 x Hr8}
                             a  =  0.22 mS/m  =   2.2jtS/cm

where:        The atomic weight of Na+  = 23, G' = 35, and the molecular weight of NaCl =
              58.

An alternative method for relating specific conductance  to solution concentration that accounts
for ionic strength is given by the following formula (Laxen, 1977):

                                  L =  1000 • 2 (X,,  • Nn)
where:
L
X,,
N
                                   specific conductance, /xS/cm
                                   equivalent conductance of ion n
                                   normality of ion n (millieq-1'1 x 10"3)
                                                                                       17.

-------
                                      Table 5
                       IONIC CONDUCTIVITY AND MOBILITY5
      Ion
Ag
Al
Ba
Be
Ca
Cd
Co
Cr
Cu
Fe
Fe
H
Hg2
Hg
K
Li
Mg
Mn
NH4
N2H5
Na
Ni
Pb
Ra
Rb
Sr
U02
Zn
Au (CN) 2
Au (CN) 4
CN
CNO
Cl
C1O2
C103
C104
F
Fe(CN)6
Fe(CN}6
H2AsO4
HC03
HF2
HS
HS03
H2SbO4
Mno4
MOO4 .
N(CN}2
N03
OCN
OH
P04
SCN
S04
Sb(OH}6
SeCN
Se04
1 +
3 +
2 +
2 +
2 +
2+
2+
3 +
2 +
2+
3 +
1 +
2 +
2+
1 +
1+
2 +
2+
1+
1+
1+
2 +
2 +
2+
1+
2+
2+
2+
1-
1-
1-
1-
1-
1-
1-
1-
1-
4-
3-
1-
1-
1-
1-
1-
1-
1-
1-
1-
1-
1-
1-
3-
1-
2-
1-
1-
2-
  lonic-X."
Conductivity
10-4 m:-mho

  61.90
  61.00
  63.60
  45.00
  59.47
  54.00
  55.00
  67.00
  53.60
  54.00
  68.00
 349.65
  68.60
  63.60
  73.48
  38.66
  53.00
  53.40
  73.50
  59.00
  50.08
  50.00
  71.00
  66.80
  77.80
  59.40
  32.00
  52.80
  50.00
  36.00
  78.00
  64.60
  76.31
  52.00
  64.60
  67.30
  55.40
 110.40
 100.90
  34.00
  44.50
  75.00
  65.00
  50.00
  31.00
  61.30
  74.50
  54.50
  71.42
  64.60
 198.00
  69.00
  66.00
  80.00
  31.90
  64.70
  75.70
 lonic-M
 Mobility
                                            6.42
                                            6.32
                                            6.59
                                            4.66
                                            6.16
                                             .60
                                             .70
                                             .94
                                             .55
                                             .60
                                             .05
                                            7.
                                            4,
                                            5.
                                            7.
 5
 5
 6
 5
 5
 7
36.24
 7.11
 6.59
  .62
  .00
 5.49
 5.53
 7.62
 6.11
 5.19
  .18
  .36
 6.92
 8.06
 6.16
 3.32
 5.47
 5.18
 3.73
 8.08
 6.65
 7.91
 5.39
 6.69
 6.97
 5.74
11.44
10.46
 3.52
 4.61
 7.77
 6.74
 5.18
 3.21
 6.35
 7.72
 5.65
 7.40
 6.69
20.52
 7.15
 6.84
 8.29
 3.31
 6.71
 7.85
  mg/L to
Milliequivalents
                                                                      Coefficient
0.00927
0.11119
0.01456
0.22190
0.04990
0.01779
0.03394
0.05770
0.03147
0.03581
0.05372
0.99216
0.00453
0.00997
0.02558
0.14407
0.08229
0.03640
0.05544
0.03025'
0.04350
0.03407
0.00965
0.00885
0.01170
0.02283
0.00741
0.03059
0.00402
0.00332
0.03844
0.02380
0.02821
0.01483
0.01198
0.01006
0.05264
0.01887
0.01415
0.00710
0.01639
0.02564
0.03024
0.01234
0.00533
0.00841
0.00625
0.01514
0.01613
0.02380
0.05880
0.03159
0.01722
0.02082
0.00447
0.00953
0.01310
44.6
217.6
119.1
102.0
115.3
110.3
111.2
229.2
109.9
110.3
231.7
110.7
123.7
119.1
47.2
39.2
109.2
109.8
47.2
43.9
41.8
106.6
126.0
122.1
48.2
115.3
90.1
109.2
41.8
38.6
48.3
45.2
47.9
42.3
45.2
45.8
43.1
527.5
299.8
38.1
40.6
47.6
45.3
41.8
37.4
44.4
47.5
42.9
46.7
45.2
75.9
233.8
45.5
134.2
37.7
45.2
130.4
      5 Table modified after Vanysek, 1993; and Laxen, 1977. Values given for standard temperature of 25'C.
18.

-------
The concentration in milligrams per liter can be converted to milliequivalents using conversion
factors listed in Table 5. The equivalent conductance for each ion is calculated using the formula
(Laxen, 1977):
                              X.  =  Xn° + [x]
where:               X,,0     =   equivalent conductance for ion n in an infinite dilute solution,
                                 cm2/(equivalent • Q) = 10"4 m2 • mho
                     [x]     =   conversion coefficient for ion n
                      I      =   ionic strength

Values of K°  are listed in Table 5. Values of coefficient [x] incorporates ionic charge and are
listed in Table 5. For a detailed discussion of the derivation of coefficient [x] see Laxen (1977).

For a dilute solution, ionic strength can be calculated from the normality using the formula
(Laxen, 1977):
                                    I = 0.5  2 (Nn • zn)

where:               Nn     =   normality of ion n
                     zn      =   valence of ion n

Using the above formulas, a computed value for the NaCl solution specific conductance of
approximately 2.2 jiS/cm is obtained.  Applying this value to the regression formula given on
page 15, a regression coefficient of 0.46 (1/2.2)  is found. This coefficient is slightly less than that
found for natural waters which usually contain some divalent ions.
Moisture Content

Rhoades and others (1976) evaluated the relationship between soil electrical conductivity (EC),
soil water content (6), soil salinity, and other pertinent soil properties. They developed a two
parallel conductor model for bulk soil electrical conductivity (ECa)..  The two conductors are due
to a bulk liquid-phase conductivity which depends linearly on the electrical conductivity of the
soil water (EQ,), and a bulk surface conductivity (ECS) from exchangeable ions at the solid-liquid
interface.  This relationships can be expressed  in the formula:
                                                                                         19.

-------
                                ECa = EC, • 6 • T + EC,

where:               6      =     volumetric water content
          v           T      —     transmission coefficient which accounts for   tortuosity of
                                  electrical current flow

An empirical formula for T is given by:

                                     T =  a-6 + b

where:               a and b are regression constants taken from a plot of
                     {(ECa - EC,)/ECw}/6 versus 6

Rhoades and others (1976) found the surface conductivity greater for finer-textured soils.  When
water content was less than a threshold value (6, = -b/a), the conductivity of the soil due to the
pore water electrolyte was zero.  Their experiment in four soils found 9, to range from 0.05 to
0.12.  Rhoades and others (1976) concluded that it may be possible to estimate EC* and
coefficients a and b based  on soil texture and mineralogy. Thus pore water electrical conductivity
(EC*) could be determined by measuring the bulk  electrical conductivity (ECa) and the
volumetric water content (6).

Keller and Frischknecht (1966) discuss the resistivity of rocks and note that pores may be filled
with electrolytes (water), gases (air), or non-conductive fluids (oil).  They give the following
formula for the relationship between the bulk resistivity and the degree of saturation:

                                P/Pioo -  S,-1   ;  Sw>Swe

where:               p      =     bulk resistivity of partially saturated rock
                     p100    =     bulk resistivity of fully saturated rock
                     nj     =     experimentally derived parameter « 2
                     S,     =     fraction of pore volume filled with electrolyte
                     Jv     =     critical water content
This relationship holds provided  the water content is above a critical value (S^) based on texture.
S^c is defined as the  point where a continuous film of water still covers the rock surfaces.  S^
varies from approximately 25 percent for sandstone to 80 percent for igneous rocks.
 20.

-------
When the moisture content drops below the critical saturation, the formula for relating resistivity
to degree of saturation is:

                               P/Pioo  =  a-Sw"2  ;   SW
-------
            III.  APPLICATION OF GEOPHYSICAL METHODS
                          TO AMD INVESTIGATIONS
                                GENERAL DISCUSSION

The following sections present reviews of geophysical methods as they relate to investigation of
AMD pollution. Each of the sections discuss the basic geophysical theory, the historic and
potential application to AMD investigations, and field procedures. More detailed discussion is
provided for those methods that have direct applicability to AMD problems: D.C. resistivity,
mise-a-la-masse, induced polarization, electromagnetics, and self potential.  Methods that do not
have a direct relationship to AMD pollution (seismic, gravity, magnetometer, ground penetrating
radar) are discussed briefly and references are provided. The final section on borehole
geophysical methods is not intended as a detailed review but as a source of references for more
detailed information.

The next section is about optimum sample spacing and the probability of finding an object, given
the sample spacing and the size of the object. The formulas are primarily for a two dimensional
space, that is, the object to be found is assumed to lie on the plane of the sampling grid. It
should be recognized  that the sampling space of geophysical methods is really three dimensional
and complex. The ability of a geophysical method to detect a buried object is related to the
sensitivity of the instrument, the basic physics of the method, the geometry of the target, the
depth of the target, and the location of the target with respect to the sampling points. Thus
selection of the appropriate geophysical method(s) for  an investigation should be based on
whether the instrument can measure to the target depth; then the issue of grid spacing should be
addressed. Careful consideration of grid spacing and probabilities can provide insight into the
density of sampling that is necessary to find a specific size  object.
                                                                                    23.

-------
                  OPTIMUM SAMPLING IN GEOPHYSICAL SURVEYS6
         k
Geophysical surveys are generally conducted along a line or grid in order to intersect features of
interest. The selection of the proper sampling density is critical to finding the object of interest
at minimal cost.  The "optimum" distance between survey lines should be such that the
geophysical anomaly is defined.  However, the concept of "optimum" should also be expressed in
terms of the probability of detecting a geophysical anomaly, since  "optimum11 implies
maximization or minimization with respect to a subjective benefit.  The "optimum11 survey grid
would always be  the one that has a 100 percent chance of finding the object of interest at the
least cost.

The issue of optimum point or line spacing has been generally addressed by Parasnis (1986) and
in more detail by Agocs (1955), as well as Kendal and Moran (1963). For geologic structures
that are uniform over a long distance, the spacing of geophysical survey lines can be kept large
provided the  individual observation points along the line are relatively close. As a general rule,
geophysical surveys can not be expected to yield information about features whose depth is much
smaller than the distance between observations (Parasnis, 1986).

The following formulas can be used to find the probability of intersecting an anomaly (Agocs,
1955):

o     For a circular anomaly of diameter D, that is less than the line spacing S, the probability
       of a line crossing the anomaly would be:

                                        P =  D/S

o     For a randomly distributed, elongated anomaly of length L, where the line spacing S is
       greater than L, the probability of crossing the anomaly would be:

                                   P - (2-L)/(7t-S)
 24.
        6References begin on page 100.

-------
o      For a randomly distributed anomaly having a length L, that is greater than the line
       spacing S, the probability of crossing the anomaly would be:


              P =  [2 • L / (TC • S)] • {1 - [(1 - S2) / L2]1*} + (2 / TC) • arccos(S / L)

                                          or

                    P = {[(2 • L) / (TC • S)] • [1 - sin(00)]} + {(2 • 60) / TC}


       where:               S/L   =     cos(60)
                           60     =     degrees in radians
o      For a randomly distributed rectangular anomaly of length L and width W, whose diagonal
       is less than the line spacing S, the probability of crossing the anomaly would be:

                               P  = 2-(L + W)/(7t-S)
o      For irregular shaped anomaly of perimeter A, the probability of crossing the anomaly
       would be:

                                    P =  A/(rt-S)
o      For a grid spacing of S and T, where T > S, and where an elongated anomaly of length
       L, L < S, the probability of crossing the anomaly would be:

                          P  =  {2-L-(S +  T)-L2}/{7t-S-T}
o      For a circular anomaly of diameter D where D is smaller than the side of the grid, the
       probability of crossing a randomly distributed anomaly would be:

                          P  - D-{(S + T)-(rc-D)}/{S-T}
                                                                                     25.

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                      D.C. RESISTIVITY GEOPHYSICAL METHODS7
GENERAL DISCUSSION

Basic Principles

Resistivity using direct current is one of the oldest surface geophysical methods. Conrad
Schlumberger first applied D.C. resistivity in 1912 (Koefoed, 1979). The D.C. resistivity method
investigates changes in subsurface layering geoelectric properties by passing direct current or low
frequency alternating current through electrodes placed in the earth and then measuring the drop
in potential (voltage) across another set of electrodes.

When resistivity measurements are made in an earth that is a homogeneous, isotropic, semi-
infinite half-space, the value measured is the true resistivity of the earth.  However, the earth is
made of layers of materials that vary laterally and vertically in resistivity.  Thus, the values of
resistivity measured with resistivity methods are apparent and not true. The measured resistivity
is a function  of electrode geometry, electrode spacing, the true resistivity and thickness of each
layer, layer dip, and lateral variations within layers.

The resistivity of a homogeneous, isotropic material between two electrodes can be found using
the formula:
 where:
P
A
L
R
AV
        ^References begin on page 104.
 26.
p  =  (A/L)-R

electrical resistivity, ii-m
cross-sectional area, m2
length of current flow path, m
AV/I, Ohm's law
difference in voltage potential
across two electrodes, volts
electrical current, amperes

-------
Geoelectric properties of a layer (i) can be described by resistivity (p;) and thickness (h;).  Using
these two parameters, other geoelectric properties can be defined (Keller and Frischknecht, 1966;
Ward, 1990; and Zohdy and others, 1974).

For a column of homogeneous earth material with a unit cross-sectional area made of layers of
differing resistivity and thicknesses, a longitudinal unit conductance (S;) for layer i taken parallel
to the layering, and a transverse unit resistance (T;) taken normal to the layering can be defined.
Geoelectric parameters T; and S; are also known  as "Dar Zarrouk" parameters and are given by
the following formulas:
                                       -  h • Pi
Electrical properties of a geoelectric layer / can also be described by the longitudinal resistivity
along bedding planes (pL) and the transverse resistivity perpendicular to bedding planes (pr).
These properties are defined by the following formulas:
                                   PL  =
                                   Pr  -
The total longitudinal conductance (Sj) and transverse resistance (TT) of a column of
geoelectrically different layers can be calculated by summing the unit longitudinal conductances
and transverse resistances using the following formulas:

                                   ST   =  2(hi/Pi)
                                   T    =
                                                                                        21.

-------
For current flowing perpendicular to the layers, the average transverse resistivity (pTa) can be
found by dividing the total transverse resistance by the total ccflumn thickness using the formula:
                                   oTa   =
where:                      2(p; • h;)      =      total transverse resistance of i layers
                            Zh;           =      total column thickness of i layers

For current flowing parallel to the layers, the average longitudinal resistivity (p^) can be found
by dividing the total column thickness by the total longitudinal conductance using he using the
formula:
where:                      Z (h; / pj)     =     total longitudinal conductances of i layers
                            Z h;          =     total column thickness of / layers

Longitudinal resistivity is always smaller than transverse resistivity with the difference being a
measure of the geoelectrical anisotropy (Keller and Frischknecht, 1966). A coefficient of
anisotropy can be defined as the square root of the ratio of the total resistivity measured in the
two principle directions, along and perpendicular to layering and is given by the formula:



An average or mean square resistivity of an anisotropic block can be given as the square root of
the product of the total horizontal and vertical resistivities and is given by the formula:

                                   Pm  =   (Pr' PL)W
Equivalence and Suppression

When interpreting multilayer soundings, T and S are sometimes all that can be determined
uniquely because various combinations of h; and p; can give the same or equivalent sounding

28.

-------
curve.  This is the problem of "equivalence" and is important when interpreting soundings
(Maillet, 1941; Parasnis, 1986; Zohdy and others, 1974; and Telford and others, 1990).  As an
example, a thin resistive layer is characterized by T = h; •  p;, whereas a thin conductive layer is
characterized by S = h; / p;. Although T and S can be found, hi and PJ cannot. Thus, it is
impossible to distinguish between two highly resistive beds of different thickness and resistivity if
the product h; • p; is the same, or between two highly conductive beds if the ratio of h; / ft is the
same (Telford and others, 1990). Also, two or more layers can be combined to provide an
equivalent sounding curve. Other methods of geophysical survey  or actual subsurface layering
information are needed to confine the solution.

Another problem with interpretation of D.C. resistivity surveys is that of suppression. When the
thickness of a layer is small compared to its depth, or when the thickness is small compared to
the layers above and below, the influence of the layer on the apparent resistivity measured at the
surface is small.  The presence of the layer is unknown or suppressed, except for layers of
extremely high or extremely low resistivity.

Many computer programs that model resistivity sounding curves for a multilayer earth can also
produce  a correlation matrix that indicates sensitivity of the relationship between h; and pj.  Ward
(1990) states that correlation coefficients must be greater  than 0.72 to be statistically significant.
Survey Methods

Electrical resistivity surveys can be performed in three general configurations:  vertical sounding,
horizontal profiling, and a combination of the two. Two other methods, mise-a-la-masse and
induced polarization, will be discussed separately.

Vertical sounding (VES) is a process by which the separation between the current and potential
electrodes is progressively increased over a fixed central point.  This causes the array to measure
the apparent resistivity at progressively deeper depths. VES is used to investigate vertical
changes in subsurface layering and produces a geoelectrical cross section.

Horizontal profiling is a method for measuring the lateral variations in resistivity by progressing
the survey along a linear traverse  using a constant electrode spacing. A constant spacing implies

                                                                                        29.

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a constant depth of investigation.  Horizontal profiling is primarily used to locate geologic
structures sjuch as buried channels, faults, dikes, and anomalou's 2-D and 3-D bodies.

The combined method repeats the horizontal profiling method at wider electrode spacings to
obtain a series of profiles that can be presented as a geoelectric pseudo-cross-section.
Array Electrode Geometries

The most common D.C. resistivity methods use four electrodes placed in a straight line with
various geometries. One pair of electrodes induces current into the subsurface, and the other
pair measures the difference in electrical potential or voltage. When two current electrodes are
placed into the earth's surface  and an external D.C. current (or with more modem instruments a
low frequency alternating current) is applied, electrical current flows through the earth from one
electrode to the other as shown in Figure 3a for a homogeneous, isotropic earth model. The
lines of current flow are perpendicular to the lines of drop in electrical potential, or the
equipotential lines. The  ideal  current lines shown in Figure 3a are marked with the percentage
of current carried by each line. The drop in potential between successive equipotential lines is
constant.

The electrode configurations commonly used in D.C. resistivity studies (Zohdy and others, 1974)
are:
                             1.  Wenner
                            2.  Schlumberger
                            3.  Dipole-Dipole
                            4.  Lee-partitioning

The Wenner array uses four electrodes placed in a straight line and spaced  at equal intervals (a)
as shown in Figures 3b and 4a. The apparent resistivity (pw) is a function of the distance between
electrodes and is given by the  formula in Figure 4a.

The Schlumberger array  uses four electrodes along a straight line but with an irregular spacing
where the distance AB between the current electrodes is equal to or greater than 5 times the
distance MN between the potential electrodes as shown in Figure 4b. The  apparent resistivity is

30.

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          SURFACE
Figure 3a.
 Figure 3b.
Potential and current distribution in a vertical plane along the line of electrodes.
Current lines of flow each carry one-tenth of the total current. The potential drop
between successive equipotential lines is constant (from van Nostrand and Cook,
1966).
                                    Current
                                     Source
                                         .Current  Meter
                                                                            Surface
                                                                    Current
                                                                    Voltage
 Wenner array D.C. resistivity equipotential and current flow lines.  Current is
 induces in outer set of electrodes, Q and Q, and voltage drop is measured across
 inner set of electrodes, Pj and P2 (from USEPA, 1986).

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a function of the distance AB/2.  In practice, MN is kept constant and AB/2 is expanded to
obtain deeper soundings. As the distance AB/2 widens, the voltage drop across MN is reduced.
At some point, it is necessary to increase the MN distance to obtain deeper soundings.  The
apparent resistivity (pB) is given by the formula in Figure 4b.

The Lee-partitioning array is similar to the Wenner array except that an additional electrode (O) is
placed midway between the M and N potential electrodes.  Potential differences are measured
between the MO and NO. The formula for calculating apparent resistivity (p,) is given in Figure 4c.

The dipole-dipole array has several configurations but the most common are the axial and
equatorial as shown in Figure 4d and e, respectively (Parasnis, 1986; and Zohdy and others,
1974). The dipole-dipole array differs from the others in that the current electrodes are
separated from the voltage electrodes. The separation between each electrode within  a pair is
often kept the same  and is significantly smaller than the distance (r) between the current and
potential electrodes as shown in Figure 4d. Dipole-dipole has an advantage over Schlumberger,
Wenner and Lee-partition because shorter AB and MN spacings are needed for deep
penetration. Disadvantages are that the dipole-dipole is harder to interpret and local lateral
variations are harder to detect. An  alternative to the dipole-dipole method is to enlarge  the
separation at one or both pairs of electrodes so that they act as bipoles creating a bipole-dipole
or bipole-bipole array. Figure 4e shows an equatorial bipole-dipole array and formulas for
calculating the apparent resistivity for the  axial (pa) and equatorial (pe) arrays.
Depth of Investigation

The Wenner and Schlumberger methods are used most often in environmental and ground-water
investigations because they are the easiest to implement and interpret. Interpretation of
resistivity measurements assumes that the earth is made of a sequence of geoelectrically distinct
layers of finite thicknesses.  Contacts between layers are assumed to be horizontal to the array
spacing.

When selecting the type of array to use for a survey, it is important to note that the actual depth
 of investigation of each type of array is governed by a number of factors including signal-to-noise
 ratio, sensitivity to surficial inhomogeneities, sensitivity to bedrock topography, sensitivity to dip

 32.

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                   M          N
a.
             • a	»4-«	a
             WENNER ELECTRODE ARRAY
                                                       Pw =
                        MM
b.       h	AB/2	4«	AB/2-
          SCHLUMBERGER ELECTRODE ARRAY
                                                              (AB/2)8—(MN/2)3
                                                       Hs = ir	

                                                                     MN
                   M    O    N
                                          B
                                          •
c.
       LEE—PARTITIONING ELECTRODE ARRAY
                                  H
          •  i   •	•   .   •
          A Q  B             M  °  N

               AXIAL OR  POLAR
                                                       P.
                                                           (AB)(MN)
                                                        Pc =
 Figure 4.
                   EQUATORIAL



             Electrode configurations for D.C. resistivity arrays with formulas for interpreting

             apparent resistivity (after Zohdy and others, 1974).
                                                                                 33V

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of layers, and others (Ward, 1990).
                                                        •
There are a number of investigations of the depth to which a surface D.C. resistivity survey can
detect a buried target or differentiate layers (Evjen, 1938; Muscat, 1941; Roy and Apparao, 1971;
Roy, 1972; Barker, 1989; and Apparao and others,  1992). One of the problems in comparing
these studies is that the definition of the depth of investigation or depth of detection is not
consistent. The depth of investigation for resistivity methods can be defined as the depth to or
below which one-half  of the total current penetrates (Eyjen, 1938).  Evjen found this depth to be
approximately equal to half the current electrode spacing for a homogeneous earth. An alternate
definition (Roy and Apparao, 1971) of the depth of investigation characteristic is the point at
which a thin horizontal layer contributes the maximum amount of the total signal  at the surface.
Roy and Apparao (1971) and Roy (1972) found for four methods the following values for the
depth of investigation based on a distance of L between the extreme electrodes:

                     Equatorial dipole-dipole     0.125-L
                     Polar dipole-dipole          0.195-L
                     Schlumberger               0.125-L
                     Wenner                    0.110-L

Barker (1989) reviewed the issue of the depth of investigation and defined a normalized depth of
investigation characteristic curve. Barker (1989) found the depth of investigation  to be 0.170-L
for the Wenner array, 0.190-L for the Schlumberger, and 0.250-L for the dipole-dipole. Where L
is the distance between extreme electrodes.

Apparao and others (1992) defined the depth of detection of an electrode array as the depth
below which a target cannot be detected assuming a minimal anomaly of 10 percent. This
minimal anomaly is based in part on the  findings of van Nostrand (1953) that the  limiting depth
of detection for an infinitely conducting buried spherical target is approximately equal to the
radius of the sphere for the Wenner array. The depth of detection is not easily predicted and  is
dependent upon the shape, size, and conductivity contrast of the  target, and orientation of the
sounding array relative to the target.

Ward (1990) presents a detailed discussion of the factors influencing the depth of investigation
and sensitivity of each array.  In general, the depth of investigation  for the Wenner,
34.

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Schlumberger and Lee-partition methods is predominantly controlled by the distance between

current electrodes. Thus equal current electrode spacing produces nearly equal depths of

investigation with the Schlumberger array having a slightly greater penetrating depth and

resolving power in VES surveys than the Wenner for the same current electrode (AB) spacing.

For horizontal profiling they are approximately equal and the dipole-dipole survey is considered

less sensitive (Ward, 1990).  Barker (1989) notes that for all methods a low resistivity layer at the

surface reduces the depth of investigation.


Today a more practical approach for evaluating the depth of investigation is to conduct several

forward modeling computer simulations to evaluate the sensitivity of the method in detecting the

assumed target(s).
Comparison of Arrays


Each D.C. resistivity method has advantages and disadvantages over the others. Zohdy and

others (1974), Pennington (1985), and Ward (1990) provide a detailed discussion of the pros and

cons of various resistivity methods.  The following is a list of items that can be used to compare

the four methods:
o      Most D.C. resistivity methods assume the subsurface layers are horizontal and
       geoelectrically homogeneous. This condition is almost never met, and small lateral
       variations in near-surface resistivity cause noise and relative error that affects Wenner,
       Schlumberger and Lee-partition methods more than the dipole-dipole method.  However,
       because the Schlumberger array maintains a constant spacing for the potential electrodes
       for several current electrode spacings, the error caused by lateral change is smaller and
       more easily detected than with  the Wenner array. As long as the lateral extent of the
       near-surface inhomogeneity is small compared to the separation of the current electrodes,
       the relative error produced by the Schlumberger array will be the same for all
       measurements with the same potential electrode spacing.

o      Dipole-dipole arrays are generally applied to deep investigations. The separation of the
       current and potential electrodes allows for shorter cables than the other methods for a
       particular depth of investigation. However, the potential difference measured with the
       dipole-dipole is smaller than with the others and requires more current input and a more
       sensitive instrument for measuring the voltage potential.
                                                                                       35.

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o      The potential electrode spacing (MN) for the Schlumberger array is always kept, small
       compared to the current electrode spacing (AB) which should be greater than or equal to
       at teast 5 times the MN spacing. With the Wenner array, the current electrode spacing is
       always equal to 3 times the potential electrode spacing.

o      Schlumberger MN spacing is changed occasionally, as the signal becomes weak, while the
       AB spacing is changed between each measurement.  With the Wenner and Lee-partition
       arrays MN is always changed along with the AB electrodes  for each measurement.

o      The AB electrode spacing for the Schlumberger, Wenner and Lee-partition arrays should
       be at least 3 times and preferably 5 to 10 times the maximum depth of interest.

o      The smallest electrode spacing for the Schlumberger, Wenner and Lee-partition arrays
       should.be less than one-half the minimum depth at which a change in material is
       expected.

o      Spacing of electrodes should be logarithmic.  That is, data should be taken at roughly
       equal spacings as plotted along a log scale of AB or AB/2 spacing.

o      Dipole-dipole can provide more details on the dip of layering than other methods.

o      Stray currents in industrial areas and telluric currents measured with long array spreads
       affect the Schlumberger array less than Wenner and Lee-partition arrays.

o      Drift or unstable potential measurements from electrodes in the ground are less with a
       Schlumberger array since the potential electrodes tend to stabilize after  5 to 10 minutes.

o      Current leakage from cables is less of a  problem with the dipole-dipole array.  The
       problem of inductive coupling is minimized.

o      Special methods of interpretation are needed for dipole-dipole.

o      It is possible to get an indication of the  lateral change in the subsurface with the Lee-
       partition array, but data reduction is not simple.

o      The Schlumberger sounding curve is discontinuous whenever the MN spacing is enlarged.
       The curve is shifted upward or downward to smooth the sounding plot.  The discontinuity
       is an indication of local inhomogeneities near the MN electrodes.
Interpretation of Resistivity Data


Today, VES is interpreted by using a numerical model to find a best-fit match of the data to an

ideal geoelectric VES curve (Ward,  1990; Zohdy and others, 1974).  Horizontal profiling is
generally interpreted qualitatively by plotting the field apparent resistivity values as either a linear

trend plotted on an X-Y graph, or by contouring a map of apparent resistivity at a specific
 36.

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electrode spacing.  The contour map presents a two-dimensional picture of the apparent
resistivity at a constant a-spacing but at an apparent depth.

The combined method of horizontal profiling at different electrode spacings usually presents the
data as a contoured geoelectric cross section or pseudo-section (Ward, 1990; Hallof, 1992).  The
resistivity contours are pseudo-contours since the field apparent resistivity data are arbitrarily
assigned to a point below the center of the array at a depth often set equal to one-half the
current electrode spacing.  Other methods of cross section interpretation use numerical computer
models to estimate the distribution, resistivity and thickness of layers (Hohmann, 1982).

Interpretation of D.C. resistivity data can be a complex task when the subsurface is not ideal, that
is, not a homogeneous, horizontally-layered and laterally extensive half-space with geometrically-
simple-shaped, homogeneous resistivity targets. Prior to the availability of computers,
interpretation of VES data was by matching field data to Master Curves for 2, 3, 4 or 5 layer
models developed based on various assumptions about the contrast in layer resistivity (Mooney
and Wetzel, 1954; Flathe, 1963; Orellana and Mooney, 1966; Orellana and Mooney, 1970; Larzeg,
1973; and Koefoed, 1979). The use of master curves is limited because the infinite combinations
of resistivity and layer thickness could  not possibly be tabulated.  Today computer simulations
have replaced curve matching.

The shape of the apparent resistivity curve is influenced by the electrode spacing, the width of
the discontinuity, the angle between the profile line and the strike of the discontinuity, and the
resistivity of the juxtaposed rock units.  There can be abrupt changes in the slope of the apparent
resistivity curve where the electrodes of a horizontal profile survey cross a buried structure.  A
structure such as a fault zone can either increase or decrease conductivity depending on the
nature of the surrounding rock and the material that fills the fault zone.

One of the basic assumptions for D.C. resistivity surveys is that they are  conducted across a
homogeneous, horizontal ground. When variations in slope are greater that 10 degrees or  highly
irregular, there are noticeable effects that should be corrected (Fox and  others,  1980; Telford and
others, 1990; and Ward, 1990).  Impacts of topography are caused by variations in depth to
layering, and moisture content as well  as concentration of current flux in valleys and divergence
of current beneath hills (Fox and others, 1980; Holcomb and Jiracek, 1984; and Ward, 1990).
                                                                                       37.

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Any of the electrode array configurations discussed above can be used for mapping lateral
geologic changes, but the profiles of each array differ considerably (Telford and others, 1990).
The variation in array profiles is more dramatic with a wide zone of anomalous resistivity, such as
a fault zone or dike, than with a contact between two rock types. Telford and others (1990)
demonstrate that for zones of anomalous resistivity, the half Schlumberger array best reproduces
the shape of the anomalous zone.

The Wenner and Schlumberger resistivity methods are' best suited for delineating horizontal
layers and vertical contacts, and are less useful for bodies of irregular shape (Ward, 1990; and
Telford and others,  1990). These resistivity methods are not particularity sensitive to 3-D
anomalies unless the depth of the anomaly is less than the radius of the anomaly.  For the
Schlumberger array, a maximum response of 12 percent contrast occurs when the depth to the
center of a buried sphere is equal to the sphere's diameter.  Hallof (1992) indicates that for
buried spherical shaped bodies the dipole-dipole  array provides the largest anomaly.

Numerous studies have been conducted to determine the response of surface D.C. resistivity to
various buried simple-geometric-shaped objects (van Nostrand, 1953; Vozoff, 1958; Unz, 1963;
Al-Chalabi, 1969; Zohdy, 1969b and 1970b; Coggon, 1971; Stefanescu and Stefanescu, 1974;
Koefoed, 1976a and 1976b; Dey and Morrison, 1979; Speigel and others, 1980; Pridmore and
others, 1981; Hohmann, 1982; Mundry, 1984; and Tripp and others, 1984; and Griffiths and
Barker, 1993).

Interpretation of resistivity soundings and profiles can now be readily done using forward and/or
inverse modeling on personal computers (Koefoed, 1979; Mooney,  1980; Interpex, 1988; Bankey
and Anderson, 1989; Orndorff and others, 1989; Zohdy and Bisdorf, 1989; and Bisdorf and
Zohdy, 1990).  However, the interpretation of data is still limited by the problems of equivalence
and suppression as discussed above.
38.

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USE OF D.C. RESISTIVITY IN AMD INVESTIGATIONS
                                                         *
          *
General Discussion

Reports on the use of conventional D.C. resistivity for investigation of AMD pollution include:
Merkel (1972),  Greenfield and Stoyer (1976), Ladwig (1982), Kehew and Groenewold (1983),
Dave and others (1986), Ebraheem and others (1990), and Benson (1993). These reports
document the application of D.C. resistivity to investigating subsurface distribution of mine
wastes, identifying possible high mineralization areas within the waste, identifying areas of
ground-water pollution and subsurface discharge, and development of empirical relationships
between apparent resistivity and ground-water quality (total  dissolved solids and specific
conductance).  The investigation methods used were VES and horizontal profiling methods with
the Wenner and Schlumberger arrays.

D.C. resistivity has been used by geologists for many years in the exploration of mineral, ground-
water and petroleum resources. Geophysical textbooks and publications on the general
application of D.C. resistivity include:  Maillet (1947), Keller and Frischknecht (1966), Kunetz
(1966), van Nostrand and Cook (1966), Morley (1967), Soil Test (1968), Bhattacharyya and Patra
(1968), Parasnis (1973 and 1986), Koefoed (1979), Telford and others (1976 and 1990), Hallof
(1992), and Keller (1993).

Publications that address the general application of surface D.C. resistivity to ground-water,
environmental and geotechnical investigations include: Flathe (1955), Breusse (1963), Krulc and
Mladenovic (1969), Zohdy and others (1974), Kelly (1976), USEPA (1978), Mooney (1980),
Griffiths and King (1981), Bruehl (1983), Dobecki and Romig (1985), Bisdorf (1985), Kelly and
others (1988), Ward (1990), Buselli and others (1992), and Henderson (1992).

Although the number of reports on D.C. resistivity investigations of AMD pollution are limited,
the literature has numerous D.C. resistivity studies on ground-water pollution.  The results of
these studies can be readily applied to investigation of AMD since they deal with geologic,
geochemical, and hydrogeologic conditions often found at mine sites.  Discussion of D.C.
resistivity investigations as applied to AMD is divided into the following general categories:
                                                                                       39.

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o      geologic  structure and stratigraphy
o      ground-water protection zones
o      aquifer properties and ground-water quality
o      fracture systems
o      landfills and waste impoundments
o      open and filled voids
o      long-term monitoring
Geologic Structure and Stratigraphy

Studies of the application of D.C. resistivity to investigation of geologic stratigraphy and
structures are numerous and include texts and reports by Logn (1954), van Nostrand and Cook
(1966), Lennox and Carlson (1967a and 1967b), Zohdy (1970), Telford and others (1976,1990),
Koefoed (1979),  Verma (1979), Ayers (1989), Barker (1990); Alfano (1993), and Kelley and
Mares (1993). Geologic structures of particular interest in AMD investigations include faults,
fracture systems, rapid lateral and vertical changes in stratigraphy, boundaries of buried waste
piles, .buried alluvial channels, buried tunnels and mine workings, open sink holes or cavities, and
filled cavities.

Application of D.C. resistivity to investigation of near vertical discontinuities is discussed in detail
by van Nostrand and Cook (1966).  Investigation of lateral discontinuities is best done using
horizontal profiling, provided the depth of the target is identified.  Selection of the proper
horizontal array should be done by evaluating existing subsurface data and conducting several
preliminary VES surveys.
                                                              •

Direct current resistivity has been used in numerous studies to investigate aquifer properties, the
extent of ground-water resources, and mapping of the depth and lateral extent of the fresh-water
salt-water interface.  These studies include: Hallenback (1953), Breusse (1963), van Dam and
Meulenkamp (1967), Flathe (1967 and 1976), Krulc and Mladenovic (1969), Zohdy (1969a),
Merkel and Kaminski (1972), Zohdy and Jackson (1973), Frohlich (1973 and 1974), van Dam
(1976), Topfer (1976), Worthington and Griffiths (1976), Worthington (1977), van Overmeeren
(1981), Urish (1983), Underwood and other (1984), Park and others (1984), Mark and others
(1986), Owen and others (1991), and Briz-Kishore (1992).

40.

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Ground-Water Protection Zones
           \
Several studies have been, conducted to evaluate the use of surface resistivity in delineating
ground-water protection zones around production wells (Mazac and others, 1987; Kalinski and
others, 1993a). These studies rely on the fact that the apparent resistivity on the surface can be
correlated with the thickness of a protective clay or silt layer once sufficient VES surveys have
been conducted.  Kalinski and others (1993b) discuss the use of D.C. resistivity soundings and
profiling to empirically estimate the vertical travel time of ground water through a protective clay
layer. Their study found that the vertical travel time can be directly proportional to the square of
the longitudinal conductance (S2) of the protective clay layer when the layer is more conductive
than the underlying aquifer, a common occurrence for fresh-water sandy aquifers overlain by a
clay layer.  Mazac and others (1987) discussed the development of ground-water protection zone
using various geophysical methods including D.C. resistivity.
Aquifer Properties and Ground-Water Quality

One of the more interesting applications of D.C. resistivity is its use in estimating physical and
chemical properties of aquifers.  Studies have demonstrated that empirical relationships can be
developed between geoelectrical parameters such as apparent resistivity, transverse resistivity,
longitudinal conductance, and apparent formation factor, and between aquifer properties such as
hydraulic conductivity, transmissivity, effective porosity, specific capacity, salinity, and total
dissolved solids (Page, 1968; Worthington, 1975 and 1976; Griffiths, 1976;  Henriet,  1976; Kelly,
1977; Mazac and others, 1979,1985,1987,1990, and 1992; Urish, 1981; Kosinski and Kelly, 1981;
Niwas and Singhal, 1981; Kelly and Reiter, 1984; Ponzini and others, 1984; Ringstad and
Bugenig, 1984; Taylor and Cherkauer, 1984;  Frohlich and Kelly, 1985; Bardossy and others, 1986;
Huntley, 1986; Ahmed and others, 1988; Mbonu and others,  1991; Ritzi and Andolsek, 1992; and
White (1994).

Typically correlations are developed by conducting VES surveys adjacent to wells where
subsurface data are known.  Researchers often correlate Dar Zarrouk parameters T and S
(transverse resistivity and longitudinal conductance) to aquifer characteristics, since they can be
determined more easily  and uniquely than layer resistivity (p;) or layer thickness (h;). Several
studies have been conducted in the attempt to remove the effects of changing pore water

                                                                                        41.

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resistivity by normalizing the apparent transverse resistivity.  They employ the use of a correction
factor (paJip,,) which is a subjective average water resistivity (p'^,) divided by the site specific pore
water resistivity (pw) (Ponzini and others, 1984; and Ahmed and others,  1988). Geostatistics have
also been used, to correlate the spatial variability of the data (Bardossy and others, 1986; and
Ahmed and others, 1988).

Niwas and  Singhal (1981) proposed the combining of Darcy's and Ohm's laws, with Dar Zarrouk
parameters to form two new fundamental laws:

                           T = K • a • R   and  T = K / (a • C)

where:               T      =      aquifer transmissivity
                     K      =      hydraulic conductivity
                     a      =      aquifer electrical conductivity = 1 / p
                     p      =      aquifer resistivity
                     R      =      transverse resistivity = h • p
                     C      =      longitudinal conductance = h / p

Niwas and Singhal propose that  if K • p is constant for a given area, the hydraulic conductivity
can be correlated to aquifer resistivity as measured by surface D.C. resistivity surveys.

In general, these researchers found that for development of empirical correlations a large number
of field sites are needed within the area of interest. The data should span the entire range of
interest, since extrapolation beyond known data is questionable.

Fracture Systems

Migration  of fluids in bedrock fractures and joints is primarily controlled by the orientation,
aperture, density and type of in-fill material.  Surface D.C. resistivity studies of fracture systems
can be used to determine the location, orientation, degree of anisotropy, direction of greatest
interconnection, fracture porosity, and transmissivity anisotropy (Merkel and Kaminski, 1972;
Habberiam, 1975; Mallik and others, 1983; Smith and Randazzo, 1989; Stewart and Wood, 1990;
Ritzi and Andolsek, 1992; Haeni and others, 1993; al Hagrey, 1994; and Lane and other, in
press).

42.

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Merkel and Kaminski (1972) conducted radial profiles centered on production wells using a
modified djpole-dipole array where one of the current electrodes was placed in the well.  They
found better delineation of the fracture systems when the well electrode was within a fracture
system.

Mallik and others (1983) found the axis of the apparent resistivity ellipse developed from radial
VES surveys compared favorably with surface-mapped fracture systems.  Also, plots of the VES
anisotropy (X) versus electrode spacing indicated variations of fracture density with depth.

Taylor and Fleming (1988) proposed a method for azimuthal resistivity surveying in jointed rock.
They used the Wenner array in radial or azimuthal resistivity surveys to determine fracture
anisotropy, direction of greatest connectivity, and fracture porosity.  They found the major axis of
the resistivity ellipse was parallel to the direction of greatest joint interconnection. When the
length of the fractures exceeds the array spacing, the major axis resistivity closely parallels the
strike of the most prominent set of fractures.  When the length of the fractures is less than the
array spacing, the major axis of resistivity is oriented in the direction of greatest connectivity for
the combined set of fractures.

Ritzi and Andolsek (1992) used the method of Taylor and Fleming to determine the orientation
of the  anisotropic transmissivity ellipse around a pumping well. The results of the azimuthal
resistivity survey compared well with the drawdown ellipse observed during a pumping well test.
They concluded that D.C. resistivity can be used to predict hydraulic characteristics of an aquifer
and can  be used to locate wells.

The basic premise of radial or azimuthal resistivity surveys is that the resistivity ellipse, a polar
plot of apparent resistivity by  azimuth, should indicate the direction through the fracture system
having the least hydraulic resistance.  The authors all noted the paradox of resistivity anisotropy
described by Keller and Frischknecht (1966):  for steeply dipping fractured or layered rock the
maximum apparent resistivity  is measured along the strike of the fracture and the minimum is
measured on the perpendicular.

Lane and others (in press), and Haeni and others (1993) used an azimuthal square-array D.C.
resistivity method to map fractures in bedrock. They found that the square-array D.C. resistivity
sounding method is more sensitive to rock anisotropy than the Schlumberger or Wenner arrays.

                                                                                         43.

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The square-array D.C. resistivity method also requires 65 percent less surface area than the
Schlumberger or Wenner arrays (Habberjam and Watkins, 1967; and Habberjam, 1972).
Landfills and Waste Impoundments

Numerous investigation of landfills and waste impoundments using D.C. resistivity methods have
been reported in the literature (Cartwright and McComas, 1968; Warner, 1969; Klefstad and
others, 1975; Stollar and Roux, 1975; USEPA, 1978; Greenhouse and Harris, 1983; Grady and
Haeni, 1984; Ruby, 1984; Stierman, 1984; Rumbaugh and others, 1987; Stierman and Ruedisili,
1988; Robert,  1989; Barker, 1990a; Buselli and others, 1990; Butler and Llopis, 1990; Carpenter
and others, 1990 and 1991; Ross and others, 1990; and al Hagrey, 1992).

The general conclusions from these waste impoundment and landfill studies regarding the
applicability of D.C. resistivity survey to AMD investigations are: 1) the lateral and vertical
extent of landfills can be defined;  2) conductive leachate plumes from landfills and waste
impoundments can be mapped provided the geoelectric contrast caused by the pollution is
sufficient to overcome natural noise and scatter in formation resistivity; 3) the integrity of landfill
cap can sometimes be assessed and fractures identified; 4) the best results are obtained when the
ground water  is shallow and the geology is relatively homogeneous.
 Open and Filled Voids

 Direct current resistivity has been used in combination with other geophysical methods such as
 seismic refraction, microgravity, and ground penetrating radar, to locate subsurface voids that are
 either filled with air, water or clay (Cook and Nostrand, 1954; Cook and Gray, 1961; Dutta and
 others, 1970; Bates, 1973; Speigel and others, 1980; Denaham and Smith, 1984; Filler and Kuo,
 1989; Smith and Randazzo, 1989; and Nelson and Haigh, 1990). These investigators generally
 found that relative to surrounding rock, air-filled cavities exhibit very high to infinite resistance
 and water- or clay-filled , low resistance.  Filler and Kuo (1989) found that D.C. resistivity
 delineated shallow limestone cavities where seismic and ground penetrating radar did not.
 Nelson and Haigh (1990) found that resistivity surveys over air-filled sinkholes were effective in
 finding caverns but not useful in locating the associated fracture systems.

 44.

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Long-Term Monitoring
          >.
The use of D.C. resistivity in conjunction with other geophysical methods for long term
monitoring of ground-water quality has been discussed by Wilt and Tsang (1985), Benson and
others (1988); Barber and others (1991); and Hanson and others, (1993). Long-term monitoring
using resistivity requires a reasonable understanding of the subsurface geology and hydrology,
knowledge of potential sources and geochemical nature of pollution, and development of
background or a baseline resistivity signature. Detection of change is complicated by seasonal
variations in soil moisture, ground-water levels,  and changes in the direction of ground-water
flow.  Monitoring highly conductive fluids such as AMD can be successful when the  contrast with
background or baseline data can be distinguished from random noise and seasonal variations.
D.C. RESISTIVITY FIELD PROCEDURES

Direct current resistivity surveys can be conducted either as vertical soundings (VES) or
horizontal profiles.  Soundings are conducted by expanding the electrode spacing over a fixed
point.  Thus VES data are collected at discrete points and provide information on the vertical
change in layer resistivity and thickness.  Horizontal profiles occur along linear or near-linear
trends, or when evaluating anisotropy along radial profiles.  The intent of a profile survey is to
observe lateral changes or to define the extent of an anomalous buried target.  Profile surveys
should be oriented to maximize the reading over the anomaly, generally perpendicular to the
strike of the target. However, special  cases exist, such as azimuthal or radial surveys, where use
of various orientations can reveal important information.

Details of field procedures for D.C. resistivity are discussed by Soil Test (1968), Zohdy (1968b),
Mooney (1980), Milsom (1989), Telford and others (1990),  and Ward (1990).  The following is a
summary of recommendations for field work::

o     For VES surveys, several sites  should be selected to develop an understanding of the
       range of resistivity and thickness.  Site selection should be based on prior knowledge of
       targets, actual subsurface data, estimates of subsurface geology, estimates of geophysical
       response based on forward modeling, and locations where information would be of
       maximum benefit. Accuracy of VES interpretations is maximized when some actual
       subsurface data are available.
                                                                                       45.

-------
o      Plot and evaluate VES survey data in the field to make adjustments in survey direction
       and spacing, or method, and to detect errors in equipment and survey procedures.
         i
o      Several VES surveys should be run perpendicular to each other to develop an
       understanding of the magnitude and impact of anisotropy.

o      VES surveys should be conducted initially in areas where horizontal profiling is planned
       in order to determine  the appropriate electrode spacing(s). Mooney (1980) recommends
       that the profiling array spacing be at least 11A to 2 times the target depth. Thus for the
       Wenner array, the overall spread would be 4 to 5 times the target depth.

o      To develop correlation between VES data and aquifer properties or ground-water quality,
       surveys should be run  adjacent to the production or monitoring wells.  If the well casing is
       conductive, the survey should be sufficiently distant enough that the well casing does not
       interfere.  Trial surveys may be needed to establish the proper distance.

o      Because horizontal profiles are conducted to find lateral changes in the subsurface,
       surveys should be run  over targets. The lateral extent of the survey should be of
       sufficient distance that the entire array is outside the target's influence. Lateral line
       spacing and  the associated probabilities of finding an target have been discussed above.

o      If profiles are repeated to obtain information at various depths, then a reduction in
       electrode spacing of between 1 to %, or an expansion in electrode spacing of 1 to 2 is
       suggested.

o      Data for VES surveys should be collected so that there are at least six data points per
       decade. One decade is equal to a factor of 10 on the logarithmic scale. To obtain this
       spacing, each new spacing is found by multiplying the previous value by 101* = 1.47.
       Increment round-off is done in practice.

o      For VES surveys the span of the data should be at least 2 decades with 2V2 to 3 decades
       preferred.

o      The small electrode spacings should be at least one-half the minimum depth at which a
       change in material is expected.

o      Resistance at the electrode should be minimized. This can be accomplished by: 1) using
       non-polarizing electrodes, such as ceramic copper-sulfate pots; 2) placing metal  electrodes
       into moist earth; 3) using two or more electrodes connected in parallel and spaced a
       meter or two apart perpendicular to the survey line; 4) pouring water  or salt water around
       the electrode.

o      Surveys should be 1 to 2 electrode spacings away from sharp changes in topography, such
       as cliffs or road cuts.  Position the survey parallel to the contour of hill or breaks in slope
       to minimize the impact of topography.

o      As the spacing of the survey is widened, there is a tendency for the value of resistance to
       decrease. However, increases or "backups" can occur when one current electrode is
       placed in or above  material of much higher resistivity, or the dip of the underlying layer is
 46.

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       changing rapidly.  A new survey should be run at right angles to evaluate the cause of
       the backup.
          >>
o      Electrodes should be seated in soils and not driven into gravels or rock.  If penetration is
       not sufficient where originally planned, offset the electrode at a right angle to the survey
      • line.
                      MISE-A-LA-MASSE GEOPHYSICAL METHODS8

The mise-a-la-masse method or charged-body potential method is a three point resistivity method.
One of the current electrodes (positive pole) is placed into the conductive mass, usually an ore
body, and the other electrode is placed sufficiently distant so that its impact is negligible. Two
options exist for the configuration of the potential electrodes. One method measures the
potential gradient by moving two electrodes placed relatively close together over the area
surrounding the buried current electrode. The other method places one potential electrode at a
reference point at a distance sufficient to be considered infinite, several hundred meters from the
study area.  The other potential electrode is then moved about to read the total potential
(Telford  and others,  1990).

The application of the mise-a-la-masse method to subsurface investigations is similar to the self-
potential method, discussed below, and is used to obtain a general understanding of the lateral
extent of a conductive body. The method is most often used in mineral exploration to determine
the extent and interconnection between pyritic ore bodies (Parasnis, 1973).  The method works
best when the conductivity of the target is 10 to 100 times that of the surrounding country rock
(Eloranta, 1984; and Beasley and Ward, 1986). Mise-a-la-masse surveys are often done in
conjunction with self-potential surveys because they can indicate the dip of the conductive body
better than  the self-potential method (Bhattacharya and others, 1984).

Papers that document the use of mise-a-la-masse in environmental and hydrogeologic studies
include: delineation of fracture systems (Jamtlid and others, 1984; and al Hagrey, 1994);
estimation of ground-water flow direction and velocity; estimation of dispersion velocity from
tracer injection tests and infiltration tests (Mazac and others, 1980,1987 and 1992; Clasen, 1988;
Cahyna, 1990; and Kelly and Mares, 1993).
        Reference begin on page 118..
                                                                                        47.

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Implementation of the mise-a-la-masse method requires that a'borehole or well be drilled into the
conductive mass or fracture system of interest. Measurements are made radially from the buried
electrode. At least two measurement points are needed on each selected azimuth in order to
determine the velocity of ground-water or plume.  These measurements should be made away
from the buried electrode at a distance of at least IVz times the  depth to the aquifer if the well is
non-conductive and three times the depth if the well is conductive (Mazac and others, 1992);
Maximum depth of exploration for ground water is approximately 10 meters with a radial
distance of approximately 150 meters (Jamtlid, 1984; Mazac and others, 1992).
                 INDUCED POLARIZATION GEOPHYSICAL METHODS9
GENERAL DISCUSSION

When direct current induced in the ground through electrodes is turned off, the drop in voltage,
as measured at two potential electrodes, is not instantaneous but decays exponentially over
several seconds to minutes of relaxation time.  This delayed voltage response is due to an
induced polarization (IP) of the earth material. The cause of the polarization is not fully
understood but is thought to be a combination of electrode polarization and membrane
polarization (Sumner, 1976;  Ward, 1990).

Electrode polarization is due in part to the presence of a diffuse double layer (DDL) of ions on
the surface of clay particles. This DDL is caused by a layer of exchangeable cations on the
surface of the negatively charged clay particles, making a double layer known as the Helmholtz
double layer (Bohn and others, 1985).  In the presence of pore water, the cations are not as
tightly held to the surface and diffuse into the aqueous phase. The concentration of the cations
decreases exponentially away from the clay surface. The thickness of the DDL is defined as the
distance over which the solution concentration is affected by the charge of the clay particles. The
thickness (d) of the DDL can be described by the formula (Ward, 1990):
       9Refermces begin on page 119.
 48.

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                              d =  K, • K • T / (2 • n • e2 • \)2)1/2
          t
where:               n      =      ion concentration of bulk solution
                     V      =      valence of ions
                     e      =      unit of electronic charge
                     K,.     =      dielectric permittivity of the fluid
                     K             Boltzmann's constant, 1.38054 x 10"16 erg/°K
                     T      =      Temperature, °K

The thickness of the DDL layer increases with an increase in Ke and T, and decreases with an
increase in n and t>.

The movement of electrical current across the DDL can be done in two ways, faradaic and non-
faradaic. Faradaic transfer is caused  by electrochemical reactions where charge is physically
carried across the interface by electron transfer. With non-faradaic transfer, no electrons are
transferred but charge is built up across the DDL. Non-faradaic transfer can be described as a
simple capacitor whose impedance varies with frequency. Electrode polarization can be
represented and predicted by a simple electrical circuit known as the Cole-Cole model of
relaxation (Ward, 1990).

A second mechanism  for polarizability is termed membrane polarization and is important in rocks
and soils with a few percent clay.  When an electric current is induced in the earth, membrane
polarization occurs within small diameter pores because the cations in the DDL block movement
of anions, but allow cations to move easily. This  creates an ion-selective membrane (Ward,
1990). The ion concentration gradient that develops opposes the flow of current, reducing ion
mobility. This reduction in ion mobility is frequency dependant, the greatest occurring at low
frequencies (0.01 Hz). As the frequency increases, the ion-selective  membrane effect becomes
less important and at  approximately 1000 Hz no longer has an affect on the mobility of ions.

Variations in the polarizability of soil and rock can be an important  phenomena in  environmental
and hydrogeologic investigations.  Polarizability can be defined as the ratio of the voltage
potential at a given time after the current is turned off,  to the potential before the  current is
turned off (Ward, 1990).  The value is expressed  as a percentage, given the term chargeability,
and derived by the formula (Ward, 1990; Kelly and Mares, 1993):

                                                                                       49.

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                                   = [AV1P(ti) / AV^] • 100%
where:               TI .,(!;)          =     apparent polarizability at time t;
                     AV,P(tj)        =     induced potential at time ti, mV
                                   =     primary potential, mV
The rate of discharge of voltage is given by:

                                   «,P  =  VrftO / VIP(t2)

where:               tx and t2 are two different times after the current is turned off

IP is similar to D.C. resistivity in that the ratio of the voltage output to the current input is a
measure of the impedance of the earth. IP surveys are conducted with electrode arrays similar to
D.C. resistivity surveys and are often done concurrently by using the value of voltage difference
at the peak of the current cycle.

IP surveys can be conducted as either vertical soundings (VES-IP), profiling or a combination of
the two. Analysis of data is similar to D.C resistivity. Computer modeling or master curves can
be used to determine vertical layering (Seigel, 1959; Dieter and others, 1969; Sadek, 1983;
Anderson and Smith, 1986; and Interpex, 1988), pseudo cross-sections can be constructed, and
contour maps of polarizability (chargeability) at equal electrode spacing can be used to interpret
subsurface features (Ward, 1990).

IP theory and array configurations, mathematical models for buried targets, are discussed by
Marshall and Madden (1959), Coggon (1973), Frische and Von Buttler (1957), Hohmann (1975
and 1988), and Fox and others (1980).

IP contrasts with D.C. resistivity in that there are two modes of data collection and
interpretation:  time domain and frequency domain. In time-domain IP, a square waveform
primary current input is turned on and off periodically.  Output voltage is measured at various
times after the  current is turned off, typically between 0.5 seconds and 2 seconds. The decay
voltage at a time tt can be normalized by the primary voltage using the formula given above for
chargeability.  The resulting units are millivolts/volt.  This normalized voltage or chargeability is a

50.

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fundamental expression of IP polarizability. The decay curve can also be integrated over a time
period (VAt) and normalized by both VRA and At to yield a unit of millivolts-millisecond/volt or
milliseconds. A third method of analyzing time domain data is the Newmont method which is a
modification of the integration method.  The standard Newmont IP cycle is 3 seconds on, 3
seconds off, and 1 second of integration time, and is typically written M331 (Ward, 1990).

The frequency-domain IP method, sometimes called complex resistivity, uses a continuous sine
wave of current as input.  Output is measured in both real (Re) and imaginary (Im) (quadrature)
components. Amplitude (pa) of the apparent resistivity and phase shift (O) of the output voltage
waveform are given by (Ward, 1990):

                                   Pa =  (Re2 + Im2)172
                                  4> = arctan(Im / Re)

In frequency-domain IP, the amplitude of the apparent resistivity and the phase shift can be
measured over several decades range in frequency (0.03 Hz to 300 Hz).  In field applications this
frequency dependence of IP polarizability becomes important because it restricts the range of
frequencies used to between 0.03 Hz and 3 Hz.  The low frequency is restricted by noise caused
by interaction of the earth's magnetic field with solar activity, while the high frequency is
restricted by electromagnetic coupling of the transmitter and receiver (Sumner, 1976; and Ward,
1990).

Results of frequency-domain IP are apparent resistivity measured at different frequencies. The
impact of frequency and polarizability can be quantified by comparing the apparent resistivity at
two frequencies. This is known as the percent frequency effect (PFE) and is given by (Ward,
1990):
                                PFE  =
where:               PI     =     apparent resistivity at frequency 1
                     p2     =     apparent resistivity at frequency 2

For normal frequency-domain IP surveys, the PFE is positive, the phase shift lags (negative phase
angle), and the secondary decay voltage (output) has the same sign as the primary voltage

                                                                                      51.

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(input). Negative IP effects are possible and commonly cause a positive phase angle.  Negative
IP effects can be caused by heterogeneities in the soil, inductive coupling, polarizability of type-K
and type-Q layer sequences, and some 2-D and 3-D bodies (Ward, 1990).

Inductive coupling has several causes including capacitive, cultural and electromagnetic (Sumner,
1976).  Capacitive coupling is caused by electrical leakage from wire-to-wire, wire-to-ground, or
electrode-to-wire. Cultural coupling is caused by artificial, grounded conductors such as pipes,
fences, or power lines.  Electromagnetic coupling is caused when the IP transmitter and receiver
circuits behave like primary and secondary windings of an electrical transformer. That is, the
primary circuit induces current into the secondary circuit. Electromagnetic coupling increases
with frequency.
USE OF IP IN AMD INVESTIGATIONS

Induced polarization has been used in conjunction with conventional D.C. resistivity studies to
explore for ore bodies, evaluate ground-water resources, investigate landfills, define limits of clay
layers, identify the fresh-water salt-water interface, and find source areas for saline ground-water
contamination (Vacquier and others, 1957; Breusse, 1963; Ogilvi, 1967; Ogilvi and Kuzmina,
1972; Bodmer and others,  1968; Cahyna, 1990, Cahyna and others, 1990; Draskovits and others,
1990; Ward, 1990; Sandberg, 1993; and Kelly and Mares, 1993).

Investigation of AMD using IP methods has not been reported in the literature, but the
environmental and hydrogeologic applications reported above suggest that IP methods can be
utilized.  The following characteristics of IP methods apply to environmental and hydrogeologic
investigations:

o      Polarizability is related to the particle size distribution. IP decay rate decreases with an
       increase in the diameter of sand particles. That is, the larger the sand, the shorter the
       relaxation time.
o      Maximum IP effects occur when the clay content ranges from 3 to 10 percent.  The lowest
       polarizability is found in clean quartz sand, and pure clay. In water saturated sandstone
       and alluvium, IP appears to be caused by clay coatings on the surface of sands and
       gravels.   .
 52.

-------
o      Flocculation and dispersion of clays can effect polarizability.  Sodium and potassium
       causes clays to disperse and swell, closing pores and reducing IP membrane effects.
       Saturation with calcium causes clays to flocculate, increasing pore interconnection and
       thereby the polarizability. For flocculated clays, polarizability is at a maximum at a clay
       content of 5 to 9 percent.

o      Polarizability of clay appears to be directly related to the cation exchange capacity.
       Montmorillonite has a higher polarizability than kaolinite.

o      Polarizability of water saturated soils is generally higher than that of for partially
       saturated soils.

o      Specific conductance of pore water can effect the polarizability. The IP decreases with an
       increase in salt concentration. IP decays more rapidly with an increase in pore water
       salinity.  At a salinity greater than 10 grams per liter, IP is negligible. IP is most strongly
       affected by cations whereas anions such as Cr, SO42', and  fOf, cause almost no change.
       IP decreases with an increase in cation valence.  For alkaline soils, the polarizability is
       greater than neutral soils.

o      Generally, polarizability decreases with decrease in resistivity. Thus clay horizons and
       saline water give smaller anomalies.

o      Polarizability is nearly constant at temperatures as high as 40°C. At greater temperatures,
       IP declines.  At temperatures below 0°C polarizability increases in coarse materials, but
       remains almost constant in clays.

o      The signal-to-noise ratio for IP is often better than for D.C. resistivity in shaley
       sandstone.

Barker (1990) used IP to investigate saline polluted ground water and  found it performed better

than D.C. resistivity for low levels of pollution in shaley sandstone.  Griffiths and others (1981)
traced the flow of NaCl pollution from an old foundry spoil pile along fissures caused by collapse

of underground mine workings. The low resistivity and high chargeability aligned with the
suspected fissures.  Chargeability increased when Cl' concentrations rose from 5 to 500 mg/1, but
decreased above 500 mg/I.


Mazac and others (1990) found that there  is no relationship between IP and aquifer hydraulic
conductivity. Cahyna and others (1990)  investigated ground-water contamination associated with

a foundry slag pile that contained heavy metals. Their laboratory studies showed the

polarizability and resistivity of the slag to be higher than that of the  surrounding country rock.

However, field studies found that only IP could delineate the slag, and that high IP chargeability

was associated with concentrated sources of ground-water pollution.
                                                                                          53.

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Although IP methods are most commonly used for mineral exploration, they can be applied to
environmental investigations including AMD.  Problems with implementation of IP surveys in
AMD studies would include:  noise caused by heterogeneities'of the waste piles, reduction in
polarizability due to increase in solution concentration, and the unknown effects of metal
hydroxide precipitates on interconnection of pores and polarizability.
                           ELECTROMAGNETIC METHODS10

GENERAL DISCUSSION

Electromagnetic (EM) geophysical surveys involve the use of frequency-domain or time-domain
electromagnetic fields to map near-surface geologic features by identifying variations in the
conductivity or resistivity of soil and rock. Frequency-domain and time-domain EM surveys can
be done on the ground, in fixed-wing aircraft or helicopters (Telford and others, 1990).
Frequency-domain EM methods include conventional very low frequency (VLF), VLF resistivity,
audio-magnetotelluric, controlled source audio-magnetotelluric, slingram, ground conductivity
meters and bore-hole methods. Time-domain EM methods include fixed-loop slingram and
central-loop configurations (McNeill, 1990). As with conventional resistivity methods, EM
methods can be used to identify and map subsurface geologic features using variable spacing for
vertical soundings or fixed spacing for cross-sectional profiles. EM methods can also be divided
into two categories based on the nature of the source wave, either planar or loop.  The planar
wave is considered to propagate across the earth's surface as a uniform, horizontal wave form.
The loop wave is generated by a local magnetic dipole. Details on the physics of EM methods
are discussed by Telford and others (1976,1990), Parasnis  (1973,1986), Sharma (1986), Dobrin
(1976), and Keller and Frischknecht (1979).

EM methods commonly used in ground water and environmental investigations include VLF,
fixed-spacing frequency-domain, ground-conductivity meters,  and center-loop  time domain.
These methods are most often used because of the commercial availability of EM systems,
portability, relatively rapid  data acquisition,  and ability to give relatively good
resolution of the subsurface. Table 6 is a summary of EM geophysical survey methods (McNeill,
1990).
         References begin on page 122.
54.

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                                         Table 6



                       Summary of Electromagnetic Survey Methods11
EM Method

Conventional
VLF








VLF
Resistivity






Audto-
Magnetotelluric
(AMT)




Controlled Source
Audio-
Magnetotelluric
(CSAMT)


Frequency Domain
Slingram







Frequency

Single









Single







Broadband






Broadband





Multiple fixed
frequency







Parameter

H









E, H,<|>







E,H






E.H.<>





B








Principal
Application
Mapping
Structures








Soundings







Soundings






Soundings





Mapping structure;
Soundings







Advantages

Very fast, One
person operation;
Inexpensive;
Works well at high
resistivity levels





Very fast, One
person operation;
Inexpensive;
Works well at high
resistivity levels



Relatively fast;
Inexpensive;
Under optimum
conditions gives
good sounding data;
Works well at high
resistivity levels
Fast;
Under optimum
conditions gives
good sounding data;
Works well at high
resistivity levels
Relatively fast for
structural mapping
and soundings;
Relatively
inexpensive;
Good measure of
bulk conductivity in
conductive ground

Disadvantages

Shallow depth;
Depth limited by
conductive surface;
Provides limited
subsurface
information;
Sensitive to
transmitter direction;
Areas of little to no
signal strength
Shallow depth;
Depth limited by
conductive surface;
Resolves at most
two layer;
Areas of little to no
signal strength;
Static shin
Signal strength
varies;
Sensitive to source
orientation;
Static shift


Expensive;
Static shirt;
Transmitter overprint



Limited ability to
sound;
Zero error;
Coil alignment and
spacing critical;
Surveys difficult in
uneven terrain;
Poor results in
resistive terrain
"Table modified after McNeill, 1990
                                                                                       55.

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                                    Table 6(continued)

                       Summary of Electromagnetic Survey Methods
EM Method

Ground
Conductivity Meter








Time-Domain EM
(TDEM)

















Frequency

Multiple fixed
frequency








Varies current
with time

















Parameter

B









dB/dt













f




Principal
Application
Mapping
conductivity and
structure;
Simple soundings






Mapping structure;
Soundings










Soundings






• Advantages

Fast for structural
and conductivity
mapping;
Relatively fast for
soundings;
Relatively
inexpensive;
Good measure of
bulk conductivity in
conductive ground
High degree of
survey flexibility;
Moderately fast;
Insensitive to
intercoil spacing and
alignment;
Zero level well
known;
Good measure of
bulk conductivity in
conductive ground;
Very fast;
Good lateral
resolution;
Good resolution of
equivalence;
Good measure of
bulk conductivity in
conductive ground
Disadvantages

Very limited
sounding data;
Shallow exploration
depth (<60m);
Non-linear response
at high conductivity;
Poor results in
resistive terrain


Relatively
expensive;
Poor results in
resistive terrain









Relatively
expensive;
Poor results in
resistive terrain


     H = Magnetic field strength
     E = Electrical field strength
     d> = Phase difference between E and H
56.

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VLF METHODS
          *
Conventional VLF

Two methods of VLF geophysical surveying are employed in environmental studies:  conventional
VLF and VLF resistivity.  Both methods rely on electromagnetic fields being transmitted by a
distant VLF transmitter. VLF transmitters are worldwide and transmit at frequencies in the band
of 15-25 kHz (kilohertz) for the purpose of military, air and marine communications (Parasnis,
1986; Telford and others,  1990). The VLF transmitter antenna is an oscillating electrical dipole
that is effectively a grounded vertical wire several hundred meters long. The electric and
magnetic fields radiate as ground, space and ionosphere waves.  At distances much greater than
the wave length, 20 kHz = 15 km, the field of the VLF dipole can be considered a uniform field
within a small area of several kilometers (km) (Telford and others, 1990). These uniform VLF
waves consist of three fields, a vertical electrical field, a horizontal magnetic field perpendicular
to the direction of the wave propagation, and a small horizontal electrical field oriented with the
direction of wave propagation.  There is no vertical magnetic field (McNeill, 1990).

Conventional VLF geophysical surveys measure components of the magnetic field. Recent
studies by McNeill and Labson  (1990) have shown that the VLF anomaly is caused by a
subsurface horizontal electrical field that induces electrical charges at the interface between two
materials of differing resistivity rather than from induced eddy currents as previously thought.
Conventional VLF response is greatest when the buried conductor strikes in the direction of the
wave propagation, and the response strength falls off with the cosine of the angle between the
conductor strike and the wave direction (McNeill,  1990).

Depth of exploration is limited to 60 to 70 percent of the skin depth, 8, as determined by the
following formula:

                                  8(m) «  500 • (p / \>)w

where:                     p      =      earth's resistivity, Q-m
                           v      =     wave frequency, Hz
                                                                                      57.

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This depth of exploration can be shallow (McNeill, 1990).  For example, for a 25 Q-m material
and a wave frequency of 15 to 25 kHz the skin depth is approximately 10 to 15 m.


Conventional VLF geophysical surveys are performed using the following steps (McNeill, 1990;

and Parasnis, 1986):


o      The presence of a conductor causes a secondary magnetic field out-of-phase with the
       primary horizontal magnetic field, creating an elliptically polarized magnetic field.

o      The VLF instrument has two solenoids aligned in the same plane, one vertical and one
       horizontal.

o      The vertical solenoid is rotated until the minimum signal is  obtained. In this orientation
       the vertical solenoid is aligned with the minor axis of the secondary ellipse of polarization
       and is perpendicular to the major axis.  The tilt angle, 6, from the vertical is equal to the
       angle between horizontal and the major axis of the ellipse of polarization.

o      The second, horizontal solenoid measures the strength of the major axis field.

o      The ratio between the signal strength of the vertical  solenoid (minor axis, b)  and the
       horizontal solenoid (major axis, a) is termed  the VLF eccentricity.

o      The tilt angle measures the in-phase or real component of the secondary vertical magnetic
       field.  While the ratio of b/a is a measure of the quadrature or imaginary component.

Although the field strength is reduced if the conductor is not aligned with the VLF wave, the tilt

angle and the b/a ratio are not significantly changed. VLF surveys have several problems that

limit their application (McNeill, 1990; and Parasnis,  1986). Conductive overburden and variations
in the thickness of overburden can generate significant VLF anomalies.  VLF is influenced by
topography and readings increase positively going uphill and negatively going downhill (Parasnis,
1986). Adjacent anomalies are superimposed and special filtering is necessary to resolve
individual responses.  In some areas of the world, VLF signal strength is too weak and a portable

transmitter is necessary.
 VLF Resistivity


 Measurement of the ratio of the amplitudes of the horizontal electric and magnetic VLF fields at

 the earth's surface can provide an apparent resistivity using the following formula (McNeill,

 1990):


 58.

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where:                     pa     =     apparent resistivity,  fl-m
                            IEI    =     amplitude of horizontal electrical field,
                                        volt/meters
                            | H |   =     amplitude of horizontal magnetic field,
                                        ampere/meter
                            p,0    =     4 • n x 10"7 Q-sec-nr1 (magnetic permeability of free
                                        space)
                           to      =     2 • jt • v
                           \)      =     wave frequency, Hz

In practice, the phase angle by which the horizontal electrical field leads the horizontal magnetic
field indicates whether resistivity increases or decreases with depth. For a homogeneous half-
space, a phase angle of 45° would be read. A greater angle indicates resistivity decreases with
depth, and a lesser angle reflects an increase in resistivity (Brooks and others, 1991).
Unfortunately, assumptions need to be made about the upper layering in order to calculate the
true resistivity (McNeill, 1990). Thus the method has difficulty resolving more than two layers.
A forward modeling computer program is available for VLF interpretation (Grantham and
Haeni, 1986).
AUDIO-MAGNETOTELLURIC AND CONTROLLED SOURCE
 AUDIO-MAGNETOTELLURIC METHODS

Two EM methods that are not usually used for environmental investigations are the audio-
magnetotelluric (AMT) and controlled source audio-magnetotelluric (CSAMT).  The AMT
method relys on naturally occurring electomagnetic fields in a frequency band of 1 Hz to 20 kHz
(lightning strikes). The CSAMT method is similar to AMT but uses a transmitter for a source
(McNeill, 1990). Both methods measure the  ratio between horizontal electrical and horizontal
magnetic fields as a function of frequency to  estimate an apparent resistivity.  The frequency
dependent apparent resistivity is modeled using a computer to estimate a simple geoelectric
section. Although the AMT method  is easy to implement, errors caused by static shift can lead
to problems with interpretation.  Static shift is caused by electrical charges that are induced at
                                                                                     59.

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the contact between layers of different resistivity which strongly amplify the local fields.  Errors
of one order of magnitude in resistivity and thickness are possible. AMT also needs a strong
signal and minimal 50/60 Hz power line noise.  The CSAMT method can be used to overcome a
weak source.  A problem arises in the need to separate the transmitter from the site of
investigation by a distance of at least three skin depths.  For a source operating at 10 Hz in a
1000 ii-m ground, a minimum separation distance of 15 km is needed. The local geoelectric
setting of the transmitter site can also overprint the investigation site readings and introduce
significant error.

Syed and others (1985) used CSAMT to locate improperly plugged oil wells that were leaking
saline brine into a shallow aquifer. To date, there is insufficient research in using CSAMT at a
geologically complex site (Bartel, 1990).
FREQUENCY-DOMAIN SLINGRAM
  AND GROUND CONDUCTIVITY METER METHODS

The EM method most often used for environmental studies, which is a variation of the slingram
method, employs the ground conductivity meter (McNeill, 1990). The slingram method consists
of using two magnetic dipole loops connected by a fixed length cable. Loops can be in any
orientation but the most common methods require that both the transmitting and receiving loops
be aligned in either horizontal (vertical dipole) or vertical (horizontal dipole). The geophysical
survey is conducted by moving the loops across the earth at a fixed interval. The coils can be
operated at multiple frequencies and separation distances to obtain either vertical soundings or
cross-sectional profiles of the conductivity of the earth's layers.
               •
The receiving loop measures the strength and phase shift of the secondary field generated by
subsurface conductors. The secondary field induced into a subsurface conductor will generally
differ in phase with the transmitted, primary field by a phase difference angle, 6. The vector of
the secondary field strength, Hs can be broken into  two components: a component parallel to
the primary field, the in-phase or real component (H$ • cos6), and a component 90° out-of-phase,
the quadrature or imaginary component (Hs • sin6) (Parasnis, 1986). To measure the in-phase
component, the coil alignment and the intercoil spacing must be carefully maintained. This is
often difficult in uneven terrain.  The vertical dipole method (horizontal loop)  is sensitive to
 60.

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narrow conductive zones or keels, and to conductive overburden whose extent is equal to or
greater than the intercoil spacing. The vertical dipole method:is also sensitive to steeply dipping,
poor conductors such as water filled fractures. The horizontal dipole method (vertical loop) is
less sensitive and is often the preferred coil orientation for geologic studies (McNeill,  1990).

Slingram systems are influenced by conductive overburden.  The primary EM field decreases by
1/e at the skin depth with a phase change of 1 radian (57.3 degrees) (Parasnis, 1986).  For
overburden thickness of less than the skin depth, the attenuation and phase change is
proportionally smaller.  Thus the phase of the field exciting the subsurface conductor is not the
same as that of the transmitter.  The secondary field from the conductor as measured by the
receiving coil is likewise influenced by the overburden.  The effect of conductive overburden is to
channel the primary field currents into the subsurface conductor where the current density is
higher than would be expected.  The channeling effect increases with increased frequency as long
as the depth of the overburden is much less then the skin depth.  At high frequencies, the
primary currents in the overburden are at the surface and screen the subsurface, eventually
causing the anomaly to disappear.

Ground conductivity meters differ from the conventional slingram system because (McNeill,
1990):
o      The low operating frequency, or low induction numbers as defined below, means that the
       receiver responses are mostly in the quadrature phase which is linearly proportional for
       low to moderate ground conductivity.
o      Ground conductivity meters are one order of magnitude more sensitive than the slingram
       systems due to operation at low induction numbers, operating frequencies low enough so
       that the skin depth is always significantly greater than the intercoil spacing.
o      The quadrature zero level is set at the factory and remains nearly constant.  This results
       in an  accurate and constant measure of the bulk ground conductivity.
o      The horizontal and vertical dipole methods give differing depth responses.  The horizontal
       dipole responds to the subsurface within approximately 75  percent of the intercoil spacing,
       while  the vertical dipole responds to the subsurface within  150 percent of the intercoil
       spacing (McNeill, 1983 and   1990).
o      The maximum intercoil spacing is 40 meters, which is less than conventional slingram
       systems.
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Two commonly used EM ground conductivity systems, EM31 and EM34-3, are produced by
Geonics Limited.  These instruments are operated at fixed intercoil spacings and frequencies and
give direct readings of apparent ground conductivity. As with the slingram systems, the receiving
coils of these instruments sense the secondary magnetic currents, H,, and compare them to the
primary field, Hp, for specific frequencies so that they operate at low-value induction numbers.
The induction number is defined as the ratio of the intercoil spacing to the skin depth (s/8).
When the induction number is much less than unity (McNeill, 1980):
                              H,/HP
where:                     H,    =     secondary magnetic field at the receiver coil
                           Hp    -     primary magnetic field at the receiver coil
                           co     =     2 • TI • i)
                           v     =     frequency, Hz
                           fi0    =     permeability of free space, Si-sec-nr1
                           c     -     ground conductivity, mho/m
                           s      =     intercoil spacing, m
                           i      =     (-!)»

Thus the apparent ground conductivity value, in mho-nr1, given by the instruments is equal to:

                              ca  =  4/(co./t0-s2-[H,/Hp]

Figures 5a and 5b show the normalized response curves, <|>v(z) and 4>h(z) and cumulative response
curves, Rv(z)and Rh(z), normalized to a "z" value that is found by dividing the depth to a layer by
the intercoil spacing. The <(>(z) function is the relative contribution to the secondary magnetic
field from a thin layer at a depth of z. The R(z) function is the relative contribution to the
secondary magnetic field from all material below a depth z, or cumulative response curve
(McNeill, 1980).
 62.

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                                                    15
                                                               20 Z
Figure 5a:     Normalized depth (z) versus relative instrument response <(>(z) curves for
              vertical and horizontal dipoles for frequency domain  EM (from McNeill,
              1980).
                  l-Op
                 05-
Figure 5b.    Normalized depth (z) versus cumulative instrument response R(z) curves
              for vertical and horizontal dipoles for frequency domain EM (from
              McNeill, 1980).
                                                                                63.

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TIME-DOMAIN EM METHODS (TDEM)
          ^
Another EM method TDEM varies the current in the primary field with time rather than
providing a continuous alternating wave (frequency-domain).  The transmitting coil is energized
with a steady electrical current.  A constant magnetic flux is inducted into the subsurface but no
secondary currents are created because the flux does not vary with time.  When the primary
current is abruptly turned-off, the primary flux falls to zero and creates transient secondary
currents within subsurface conductors because the flux is varying with time. The time dependent
decay of the magnetic field induces transient electromotive forces into the receiving coil
(McNeill, 1990; Parasnis, 1986).  Environmental studies employ two methods of TDEM: the
slingram and the central loop.

The slingram method is similar to the frequency-domain methods. That is, small transmitter and
receiver loops are moved along at a fixed separation and a cross-sectional profile of the earth is
obtained. Vertical sounding can be made by varying the intercoil separation. The TDEM
slingram method has three important differences from the frequency-domain slingram method
(McNeill, 1990):

o      TDEM is less efficient, therefore the transmitting loop must be large or multi-turn in
        order to create a large dipole moment. Common dimensions are 5 m square.
o      Constraints of the precision of intercoil spacing and alignment are less for TDEM
        because measurements are made while the transmitter is  off.
o      Zero level of the receiver is more accurate with TDEM because measurements are made
        while the transmitter is off.

The second TDEM method uses a large transmitting loop, typically  20 to 150 meters square, with
 a small receiving loop in the center of the transmitting loop or at the outer edge.  Placement of
 the receiving coil at a distance equal to the side length of the transmitter loop can reduce the
 magnitude of the inaccuracies caused by induced polarization (IP) effects. When the primary
 field is turned off, horizontal eddy currents are instantly generated near the transmitting loop that
 try to maintain the magnetic field at the same strength as before the current was turned off.
 These eddy currents increase in depth and expand radially with time. The decay of the magnetic
 field is a measure of the resistivity of the subsurface as a function of depth. TDEM has the
. following advantages over conventional D.C. resistivity for deep  sounding (McNeill,1990):
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o      TDEM is faster to cany out. The rate of surveys can be increased by laying out several
       transmitting loops with the transmitter, receiver and receiver loop moved along to collect
       data.
o      Depth of exploration of TDEM sounding is larger than the array dimension unlike D.C.
     •  resistivity which must have an array length several times the depth of interest.
o      The central loop method is insensitive to overburden conductivity variations and local
       inhomogeneities.
o      The problem of equivalence is less for TDEM than for D.C. resistivity. However, D.C.
       resistivity is superior to TDEM at resolving the resistivity and thickness of intermediate
       resistive layers.
o      Depth of exploration of the TDEM method is greater than that of the slingram method
       and can be up to several kilometers.
USE OF EM METHODS IN AMD INVESTIGATIONS

EM methods have been used for the past 20 years for evaluating subsurface geology, subsurface
hydrology and water quality. Although EM methods provide information on the subsurface
geoelectric layering that is often less precise than D.C. resistivity data, the ease of use and the
rapidity of obtaining data make it a preferred method for shallow environmental studies,
especially  for reconnaissance investigations. In recent years numerous environmental studies
have used EM methods.  Geonics Ltd. (1992a, 1992b) has published two extensive bibliographies
of studies and reports on EM geophysical methods used for:

       o      agriculture soils mapping;
       o      archaeology investigations;
       o      mapping underground voids in limestone;
       o      locating underground tanks;
       o      defining depth and extent of permafrost zones;
       o      predicting the need for cathodic protection;
       o      landfill siting and  monitoring;
       o      mapping sites of disposal of manufacturing waste,
              radioactive waste, mining waste, and petrochemical waste;
       o      mapping higher permeability aquifer channels;
       o      mapping the extent of contamination in aquifers;
                                                                                       65.

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       o      mapping anisotropy in the subsurface;
       o  ^    defining the extent of weathering, fracturing an'd faulting;
       o      identifying areas of higher ground-water yield;
       o      mapping the saltwater-fresh water interface;
       o      defining the extent of protective clay layers;
       o      exploration for mineral deposits; and
       o      exploration for petroleum reserves

While the studies listed by Geonics  do not all specifically address AMD problems, many of them
evaluate geologic and ground-water settings that are similar to those found at mine sites. The
two Geonics bibliographies are a valuable resource and can be obtained from Geonics Ltd. in
Mississauga, Ontario, Canada, at telephone (905) 670-9580, fax (905) 670-9204.

Published reports of EM investigations of AMD problems are limited and most studies were
conducted in coal mine areas rather than the pyritic hard rock mines of the western United
States.  Greenfield and Stoyer (1976) conducted one of the earliest studies. They ran four
profiles in a Pennsylvania strip mine across areas of potentially high AMD ground-water flow
using horizontal and vertical dipole EM at 40 and 60 meter spacings.  They also ran D.C.
resistivity  profiles for more detailed layering information.  Their study identified a fracture zone
that produced a high EM anomaly  due to increased saturation by high specific conductance AMD
(« 330 mmhos/m). They also noted that the same fracture system might produce a low EM
anomaly during a drought because  of more rapid drainage.

Ladwig (1982) conducted EM studies at three reclaimed coal mines that ranged in surface area
from 15 to 37 acres (6 to 15 hectares).  Chemical concentrations of the AMD at the three sites
ranged  from 20 to 4000 mg/L acidity, 18 to 800 mg/L iron and 200 to 1800 mg/L sulfate. EM
profiles using 10- and 20-meter intercoil spacing were able to identify areas of buried refuse with
high pyrite content, areas of shallow AMD ground water, source areas for surface seeps, a
subsurface drain, and delineate the previously unknown extent of the surface mine excavation.

Brooks and others (1991) used a ground conductivity meter (EM34) at a 10-meter intercoil
spacing and VLF resistivity (EM16R) over an abandoned Indiana underground and surface  coal
mine.  Specific conductance of the  acidic ground water ranged from 3,000 to 18,000 jimhos/cm.
Results of the EM surveys were combined with drill hole logs and monitoring well water quality

66.

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data to infer sources of ground-water contamination, and to delineate the hydrologic connection
between w^ste rock with a high percentage of pyrite (4 to 10 percent) and the surrounding
undisturbed materials. A high conductivity anomaly near an adjacent stream was inferred to
indicate a connection between the stream alluvium and the underground mine workings.

More recently, studies of geophysics application to AMD problems have been conducted by King
and Pesowski (1993,1991), and King and Sartorelli (1991). These studies found that VLF
resistivity and EM ground  conductivity meter surveys could delineate conductive plumes leaking
from drilling fluid disposal pits, an abandoned uranium mine, and from brine and tailing storage
facilities.  They also used EM methods to delineate buried channels beneath a proposed tailings
storage facility.

To further evaluate the potential application of EM investigations to solving AMD problems,
other studies of conductive ground water pollution with EM were reviewed.  The studies of
interest to AMD investigations can be classified into three general categories: 1) subsurface
geologic interpretation; 2) mapping saline brine; and 3) monitoring and mapping leachate from
landfills.

The use of EM methods to interpret the subsurface geology relies on a correlation between
geologic layering and conductivity contrasts.  In natural geologic materials, conductivity contrasts
occur due to variations in  porosity, permeability, grain size and pore water quality (see discussion
of geochemical and geophysical relationships, page 6). At a mine site where AMD pollution is
present, the ability to use conductivity contrasts to delineate subsurface geology can be restricted
by the random noise caused by heterogeneities of the waste piles, migration of very high specific
conductance AMD fluids into different geologic units (which masks the natural differences), a
lack of knowledge about man-made subsurface conditions such as the depth to a mine excavation
or waste fill, and the lack  of knowledge about the location of subsurface openings.

Nevertheless, EM geophysical methods have successfully delineated the general nature of
subsurface layering in several studies including:  Morgenstern and Syverson (1988), Sanders and
Cox (1988), Monier-Williams and others (1990), Mazac and others (1990), Lawrence and
Boutwell (1990), and Wrightman and others (1992). The use of EM for investigating  saline fluids
or fluids with high total dissolved solids is discussed by: Slaine and  Greenhouse (1982), Barlow
and Ryan (1985), Stewart and Brentnall (1986), Lyverse  (1989), Stewart (1990), Sartorelli and

                                                                                        67.

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others (1990), Goldstein and others (1990), Barker (1990), Street and Engel (1990), Lahti and
Hoekstra (1991), Hoekstra and others (1992), and King and Fesowski (1993). The use of EM for
         v                                             l
investigating and monitoring landfills is discussed by: Slaine and Greenhouse (1982), Greenhouse
and Harris (1983), Lasky (1985), Greenhouse and Slaine (1986), Jansen and others (1992), and al
Hagrey (1992).

Interpretation of EM data for environmental studies is often done by comparing line plots of EM
readings taken at different times or along a traverse, contouring a plan view of data from a
surveyed grid, or conducting forward and/or inverse 1-D or 2-D modeling to create geoelectric
cross sections. 'Greenhouse and Slaine (1982, 1986) recommended that EM data be normalized
to either an arbitrary constant background value or a value taken at the initial sampling (time =
0), and that the results be expressed in terms of decibels (dB) using the following formula:
                              ^decibel = 20 • 10gu(0v
where:               a^     =     apparent ground conductivity at a point x,y
                                  background apparent ground conductivity
Normalization of the conductivity has the following three advantages over the use of actual
values:

o      Normalization to a zero background puts all instrument readings on a common format
       and allows comparisons of data between instruments of differing sensitivity.
o      Logarithmic contour lines do not cluster around contaminant sources to the degree that
       linear data do.
o      The procedure is objective except for the selection of the constant background value.

By expressing the apparent conductivity in decibels it is often easier to view contrasts and
changes over background or initial values. A conductivity value of zero dB indicates no change
over background, a value of +6 dB is equal to twice background, and a value of -6 dB is equal to
one-half background. Slaine and Greenhouse (1982) suggest that a noise level of approximately
4 dB (or 1.6 times background) should be anticipated.  Greenhouse and Slaine (1986) propose
the development of a "standard section" on which a 1-D sensitivity model is run to evaluate the
impact of variations in layer thickness and layer conductivity. Based on the model results,  a
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normalized background, uncontaminated section can be selected. Greenhouse and Slaine (1986)
introduce die concept of a "formation ratio" and a "surface ratio" which are the ratios of the
conductivity of individual layers and of the whole section to the background section.  They
suggest that a formation ratio of at least 2 times background is needed to create a surface ratio
of 1.5 times background. Also, a threshold value of as much as 6 times the background is needed
to provide reliable evidence of contamination at a single station.  Greenhouse and Slaine (1986)
noted that contoured data can provide a more reliable indication of contamination over plots of
single station data.

Monier-Williams and others (1990) found that topography had a significant influence on EM
surveys of landfills due to a reduction in the depths to ground water and an underlying clay layer
as the surface elevation decreased. Adopting the decibel normalization method of Greenhouse
and Slaine, they suggested an empirical method for correcting the effects of topography.  In
effect a function, a(h), relates the surface elevation to the apparent conductivity of
uncontaminated ground. The function a(h) is then substituted for abackgrouild in the Greenhouse
and Slaine formula.

An alternative method for removing the effects of topography is to calculate the apparent
conductivity at the water table. Using EM data  and depth to the water table measured at two
wells, Emilsson and Wroblewski (1988) simultaneously solved McNeilTs 2-layer case equations to
find the conductivity at the water table.  For more complex layering, forward  computer modeling
can be used to predict the impacts of topography and overlying layer thickness. Contouring of
the depth, thickness, or the Dar Zarrouk parameters of a geoelecrric layer based  on forward or
inverse computer modeling (Interpex, 1988; and  Grantham and others, 1987) can provide a better
definition of anomalies because the noise from other layers is removed.

Lawrence and Boutwell (1990) have suggested that when detailed subsurface data are available
from drill holes and monitoring wells a multivariate regression analysis can  develop  a correlation
between EM signature and geologic layering. They recommend that a minimum of 5  data points
be used to develop such correlations.

Mazac and others (1990) suggest a method for site-specific correlation of hydraulic conductivity
and surface geoelectrical measurements for both the saturated and unsaturated zones for porous
and fractured rock.  The relationship between hydraulic conductivity, K, and resistivity, p, is non-

                                                                                       69.

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linear, but can approach linear for some parameters within a limited range. They propose a
method by which observed or interpreted parameters are placed into subclasses, each with a
specific weighting factor.  The sum of the weights pertinent to all parameters, expressed as a
percent, gives an integral indicator value, I. The integral indicator value can be contoured and a
linear relationship between saturated K and I can then be found. Once developed for a specific
site, the K-I relationship can be used to delineate areas of differing hydraulic conductivity.
EM FIELD PROCEDURES

EM surveys are conducted similarly to other surface geophysical surveys.  That is, the assumed
target area(s) is (are) identified; the objectives of the survey are defined; based on sensitivity and
applicability the proper EM methods are selected; and the survey points and grid are laid out.
EM surveys at mines often focus on known or observed sources of AMD seepage rather than the
entire site. EM surveys for AMD at mine sites can be complicated by a lack of knowledge about
historic operations, by abandoned and buried metal objects, and by the roughness of the terrain.
The following are general procedures recommended for EM surveys (McNeill, 1980, 1985b, and
1990; Geonics, 1988).  For specifics on instrument calibration and operation the operating
manual should be consulted (Geonics, 1985a and 1991).

o      Select target area and determine orientation of feature of interest, if possible.
o      Optimize sample point and line  spacing, intercoil separation based on results of test lines,
       knowledge of target dimensions and depth.
o      Determine whether ground conductivity contrast is sufficient to identify the plume or
       buried object of interest. Use knowledge of uncontaminated materials and estimates  of
       contaminant concentration and  specific conductance for a preliminary model instrument
       response.
o      Target should have an apparent ground conductivity that is at least 150 percent above or
       below background.
o      Lay out grid and identify sample points with non-conductive  stakes or markers.
o       Calibrate instrument at site as required by manufacturer.
o       Orient traverses perpendicular to strike of target, if possible.
 70.

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o      Extend traverse lines into areas of no contamination, if possible.  For point monitoring,
       select stations within and outside of the plume. Select; a background site to obtain a
       normalization value.

o      Avoid areas of irregular topography, if possible, because variations from the ideal half-
       space model can introduce error. Run survey lines along contour of slopes, if necessary.

o      To determine the degree of anisotropy, take readings at a station at differing azimuth
       orientations to find anisotropic ellipse.

o      Determine extent of interference from fences, powerlines, and pipes by making a traverse
       perpendicular any such features until EM readings stabilize.

o      EM horizontal dipole (vertical coils) is sensitive to:
              o      near surface variation in conductivity;
              o      overhead powerlines, keep receiver furthest from powerline;
              o      fences, buildings and tanks, experiment to find safe distance; and
              o      near surface metal debris.

o      EM vertical dipole (horizontal coils) is sensitive to:
              o      subsurface variations in conductivity;
              o      lithologic changes from faults;
              o      pipe lines;
              o      vertical dikes; and
              o      shallow buried metal debris.

o      Data collection using written notes and data loggers can significantly speed up field survey
       time and data processing.

o      Written notes are essential. Data loggers can introduce error when station location or
       instrument settings are not set properly.  Errors are recoverable provided written notes
       document starting and ending locations and station numbers, line orientations, intercoil
       spacings, station  spacings, field personnel, dipole orientation, and other critical survey
       information.
                           SELF POTENTIAL METHODS (SP)12


GENERAL DISCUSSION


Self potential (SP) voltages occur in the subsurface due to natural electrochemical, electrokinetic

or thermoelectric reactions.  These potentials are associated with weathering of sulfide mineral

bodies, variations in mineral content along geologic contacts, bioelectric activity, corrosion, and
         References begat on page 127.
                                                                                       71.

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thermal or pressure gradients in fluids.  There are four principal mechanisms that produce self
potentials:^) electrokinetic or streaming; 2) diffusion or liquid" junction; 3) Nerst or shale; and 4)
mineralization (Telford and others, 1990).

Electrokinetic potential, also called electrofiltration, and streaming or zeta potential, is caused
when an electrolyte flows through a capillary or more accurately a porous media. The result is a
potential difference between the ends of the passageway that can be expressed by the formula
(Parasnis, 1986):

                             Ek = -{£-p-C-P}/{4-TT-/i}

where:               £      =      dielectric constant of electrolyte, Fnr1
                     p      =      resistivity of electrolyte, ii-m
                     £      =      streaming potential, V
                     /x      =      dynamic viscosity, Pa-sec
                     P      =      pressure gradient, Pa-nr1

Dielectric constant is a measure  of the  electric polarization that occurs when an electric field is
applied and varies inversely with frequency (Telford and other, 1990). The dielectric constant is
the ratio of the specific capacity of the material to the specific capacity of a vacuum (8.85 x 1042
farads per meter, Keller, 1989).  Dielectric constant of a rock or soil material changes with water
content because water has a high dielectric constant, approximately 80 F-nv1 (see Tables 6 and 7
in Keller, 1989; and Table 5.5 in Telford and others, 1990).

Streaming potential can be found associated with flow of water through porous media.  The
streaming potential is a double layer potential between the solid and solution phase (Telford and
others, 1990). Streaming potentials are generally of minor importance in mineral exploration
surveys, but can be found associated with large negative anomalies on topographic highs and are
important in spontaneous potential borehole logging where drilling fluids flow into porous
formations.
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The diffusion potential is due to a difference between the concentrations of electrolytes and can
be expressed by the formula in millivolts (mV) (Telford et al./'1990):

                   Ed  =  - {R • 6 •  (I. - Ic) • ln(Q / Q)} / {F • TI • (Ia + Ic)}

       where         R     =     the gas constant, 8.314xlQ-3 ia-°Cl-tK
                     F      =     Faraday's constant, 96.445 kJ-V^-mor1
                     9      =     absolute temperature, °K
                     T|      =     ion valence
                     Ia     =     mobility of anions, m-s'W-nr1
                     I,.     =     mobility of cations, m-s'W-nr1
                     Q, Q =     concentration of electrolytes

The Nerst potential occurs when the  electrolyte concentrations around two identical metal
electrodes are different.  The potential can be expressed in millivolts (Telford and others, 1990)
by:

                            En =  - {R • 6 • ln(Q / Q) }/ {F • r\}

Where the coefficients are the same  as above.

For example a NaCl solution where Ia / Ij  = 1.49 produces a diffusion potential at 25 °C of
Ed -  -11.6 • log (Q / C2) and a Nerst potential of En = -59.1 • log(Q / Q) (Telford and
others, 1990).

The electrochemical self potential is  the sum of the diffusion and Nerst potentials and can be
expressed in millivolts as:

                         Ec .  - 70.7 • [(T + 273)/273) • log(Q / Q)

where:                      T     =     temperature, °K

Mineralization potential, or sulfide potential, occurs when two dissimilar metal electrodes are
immersed in a homogeneous electrolyte.  This potential is commonly found with ores containing
                                                                                        73.

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pyrite and chalcopyrite.  Although the exact mechanism for large negative anomalies over sulfide
ores is not \fully understood, the use of self-potential surveys to locate ore bodies is a common
practice in mineral exploration.  The combined effect of the diffusion and Nerst and
mineralization potentials is thought to be the principle cause of the large anomalies found with
certain mineral ores.  Theories on the causes for self potential voltages associated with mineral
deposits are discussed by Sato and Mooney (1960), Kilty (1984), Cony (1985), and Fumess (1992
and 1993).  Theories on causes for self-potential voltages associated with geothermal systems are
discussed by Corwin and Hoover (1979) and Sill (1983).
USE OF SP IN AMD INVESTIGATIONS

The use of self potential methods for investigating AMD problems is limited to areas where there
is sufficient ground-water flow to develop streaming potential anomalies because the diffusion,
Nerst and mineralization potentials are all reduced significantly by high specific conductance
electrolytes (Bogoslovsky and Ogilvy 1973,1972; and Ogilvy, Ayed and Bogoslovsky, 1969).  A
review of literature found  only two published reports on the use of SP for evaluating acidic fluids
(Stierman, 1984; and Reznik, 1990). The results of the Stierman study at an acid waste disposal
pond were too noisy to provide any definitive results.  Reznik's (1990) study at a partially
reclaimed coal mine had better results. Reznik was able to identify ground-water flow direction
and sources of AMD seepage.  Even though the application is limited SP can be useful in AMD
investigations because  it can:  1) identify zones of increased infiltration; 2) identify areas of
leakage or seepage from canals, dams, springs and reservoirs; 3) evaluate effectiveness of ground-
water drainage structures; and 4) evaluate effects  of ground-water pumping.

Studies of SP phenomena related to infiltration have been reported by Ernstson and Scherer
(1986), Bogoslovsky and Ogilvy (1970,1972,1973), and Erchul (1986).  In general, infiltration of
surface water into fractured rock or karst sink holes produces SP minimum when compared to
the surrounding lands.  This reduction in SP voltage is thought to be linearly related to the
pressure gradient as long  as the flow is laminar and not turbulent (Ogilvy, Ayed and Bogoslovsky,
 1969). Widening of fractures tends to decrease amplitude of the negative anomaly. Infilling of
fractures with sand up to  approximately 40 percent causes the negative SP  anomaly to increase.
 Infilling with clay decreases the amplitude of the anomaly.
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Studies of SP anomalies associated with canal, dam and reservoir leakage have been reported by
Ernstson and Scherer (1986), Ogilvy and Bogoslovsky (1979), JBogoslovsky and Ogilvy (1972), and
Ogilvy, Ayed and Bogoslovsky (1969).  These studies found that a negative SP anomaly is
generally associated with areas of seepage with the intensity of the anomaly increasing at higher
rates of seepage.

Bogoslovsky and Ogilvy (1973) found that for subsurface drains, positive SP anomaly
equipotential lines parallel the discharge into the drain and negative anomalies are often
associated with the drain outlet.  Bogoslovsky and Ogilvy (1973) used a SP survey to qualitatively
evaluate the effectiveness of a drainage system installed to dewater a landslide.  The SP survey
found an increase in SP values adjacent to a section of the slide that continued to fail, suggesting
that the drain was not functioning.

Schiavone and Quarto (1984) found that SP anomalies can be associated with upward flow of
fresh water across geologic discontinuities or lithologic boundaries.  Asymmetrical anomalies are
often associated with vertical boundaries while symmetrical anomalies occur over horizontal
boundaries (Schiavone and Quarto, 1984; and Fitterman, 1979). Radial flow  to a pumping well
can produce a circular positive anomaly and injection of water a negative anomaly (Schiavone
and Quarto, 1984; and Bogoslovsky and Ogilvy, 1973).

Analysis of SP survey data is generally qualitative rather than quantitative because the SP
phenomena involve several different potentials, and because the exact nature of the
mineralization potential is still uncertain.  Several authors have presented models for SP
interpretation of anomalies over vertical boundaries (Fitterman, 1979), vertical dikes (Fitterman,
1983), two-dimensional sheet-like and cylindrical sources (Satyanarayana Murty and Haricharan,
1985), spherical and elliptical ore bodies (Becker and Telford, 1965; and Telford and others,
1990), a thermal point source with vertical contact  and overburden  (Corwin and Hoover, 1979;
and Sill, 1983), current dipoles in half-space (Kilty, 1984), and surface and vertical profiles over
vertical and dipping sheet-like bodies (Furness, 1992).

Telford  and others (1990) report the depth of an ore body is approximately one-half of the  width
of the negative anomaly measured at half the amplitude, as a  rule of thumb.  Based on a limited
survey of SP studies, this rule, of thumb yields depths that are  ±100 percent of the true depth.
Data from SP surveys are generally used as qualitative information  along with other more

                                                                                       75.

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quantitative methods, such as D.C. resistivity and IP.  Data are presented either as a plan map

with contours of SP voltages or as x-y plots of SP voltage versus distance along a traverse.
SP FIELD PROCEDURES


Field procedures for SP surveys use two electrode configurations: 1) two electrodes separated by

a constant distance are moved along a survey line, using a voltmeter to measure the gradient
between them, or 2) one electrode is placed as a base point and the other electrode is moved

from station to station using a long wire. The second method has an arbitrary referenced

"absolute" SP as a base point. The base point electrode is connected to the negative terminal of
the voltmeter. When the base point needs to be moved, the potential of the new base point is
measured relative to the previous base point so that data can be corrected to a fixed reference
voltage.


Detailed  description of field procedures, precautions and instrument specifications are given by

Corwin (1990), Corwin and  Hoover (1979), and Cony (1985).  A summary of field procedures

and sources of noise is given below.
o      Electrodes for the SP survey should be non-polarizing. Standard SP electrodes are
       unglazed, porous porcelain pots with a metal terminal that is submerged in a salt solution
       of similar composition, such as copper terminals in copper sulfate solution. The
       permeable pot slowly leaks solution into the soil to make a good electrical contact with
       the ground.

o      Although porcelain electrodes are considered "non-polarizing" they do respond to such
       factors as temperature, soil moisture, and soil chemistry.  For example, for a copper-
       sulfate electrode, changes in moisture content of 1 percent can cause a +0.3 to +1.0 mV
       variation. Thus observation of soil moisture conditions can be critical when SP anomalies
       are a few tens of millivolts (Corwin, 1990).

o      Electrodes should be buried at least 6 inches for detailed surveys where the SP anomaly
       has a low amplitude, such as over geothermal systems. Random noise of sufficient
       amplitude to mask the anomaly is caused by moisture differences.

o      For detailed low amplitude SP  anomaly surveys, the ground  surrounding the electrodes
       should not be wetted and the contact resistance should be less than 50 k H. Voltage
       should be read before contact resistance because the ohm-meter will temporarily polarize
       the electrode. "
76.

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o      Temperature variations between the electrodes should be kept to a minimum to reduce
       voltage drift which can range from +0.5 to +1.0 mV per degree Centigrade.

o      A high impedance (> 1 x 107 ohm-meters) digital voltage meter should be used for field
       readings. The negative terminal of the voltmeter should be connected to the base station.

o      Base station electrode should be placed outside an area of steep SP gradients for
       maximum anomaly amplitude.

o      Base station electrode should be placed outside a chemically reduced environment such as
       a bog or marsh for maximum anomaly amplitude.

o      Power stations, pipelines with cathodic protection, culverts, grounded fences, drill-hole
       casings- and buried metallic objects should be avoided.

o      Read differences between the electrodes at least at the beginning and at the end of the
       survey to allow for correction of drift.  At the beginning of the survey the difference
       should be less than 2 mV.  If greater, the electrodes should be cleaned and filled with
       fresh electrolyte.

o      Long period telluric currents caused by temporal variation in the earth's magnetic field
       can be several hundred mV/km, and can increase the noise for a survey more than 1 km
       long.

o      Vegetated areas should be avoided because bioelectrical activity can generate SP
       anomalies as high as 150 mV/m.
                          SEISMIC GEOPHYSICAL METHODS13


Seismic geophysical methods for subsurface investigations are divided into two categories,

refraction and reflection. These methods differ in the progression of seismic energy through the

earth's layers, and in data collection and interpretation.  For shallow engineering, environmental

and hydrogeologic investigations, the seismic reflection method is not commonly used.  A brief

discussion of seismic reflection applied to environmental and hydrogeologic investigations and a

limited reference list are presented at the end of this section. The following discussion is a brief

review of seismic refraction methods as applied to AMD, environmental and hydrogeologic

problems. Haeni (1988) provides a more detailed discussion of seismic refraction applications to

hydrologic studies including an extensive annotated bibliography.
         References begin on page 130.
                                                                                       77.

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Seismic energy travels through the ground as four types of waves: 1) a compressional or
longitudinal wave (ground motion parallel to the direction of propagation); 2) a shear or
transverse wave (ground motion perpendicular to the direction of propagation); 3) a Rayleigh
wave (travels along the surface); and 4) a Love wave (travels along the surface). For engineering,
environmental, and hydrogeologic seismic surveys, the compressionalwave is utilized most often.
The shear wave is used for special engineering studies.  Surface waves are not used.
Compressional waves travel the fastest, thus they are the first to arrive at the geophones and are
the easiest to identify on the seismographic record.

Figure 6 shows schematically the ray-paths of energy from a surface source for a two-layer earth
model.  The two-layer earth model assumes ideal horizontal, homogeneous, and isotropic layers,
and velocity increasing with each deeper layer. Seismic waves travel outward from the energy
source along four ray-paths:

       o     as a direct ray-path along the surface;
       o     as a totally reflected  ray-path that strikes the interface between two layers at an
              angle greater than a  critical angle of incidence (i,.), with all of the energy reflected
              toward the surface at the velocity of the upper layer (Vx);
       o     at a critical angle of  incidence (ic) where part of the energy is reflected at velocity
              Y! and part of the energy is refracted along the interface between the two layers
              at the velocity of the deeper layer (V2); and
       o     as a ray-path that strikes the interface at an angle less than the critical angle
              where part of the energy is reflected and part is refracted downward into the
              deeper layer.

The seismic refraction method uses  the first and  third ray-path properties and the critical angle of
incidence to develop formulas for calculating layer velocity and thickness.  The critical angle of
incidence at the interface of two layers is given by:

                                    ic  =  arcsin(Vi / V2)

where:                      Vj     =     the velocity of the upper layer
                            V2     =     the velocity of the lower layer
78.

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               ENERGY SOURCE

                     S
           GEOPHONE

               G
                                                             /-•-RAY PATH OF HEAD WAVE
                                                                         "l >'c>'|
                              (3) PATH OF RAY REFRACTED
                                 AlONG V, -VjCONTACT
Figure 6.      Schematic ray-path diagram for seismic energy generated at source S and received
              at geophone G (from Zohdy and others, 1974).
   • Shotpoiml
                   Shotpoim 2
                                1    2  3  4   E  6
 /™™ snotpoint 3           ^
/7  t  9   10 VI 12    /
                                                                       Shotpoint 4     Shotpoint C

X
K
Q.
LU


!V VY| \~7T//! A '
\ \~* F ~ ~ ~ "V //
\ y / /
V / *- •— — ~ 	 MM*"!"''

Layer 1
Layer 2

Layer 3
                                              DISTANCE
 Figure 7.      Field setup of shotpoints and geophones for delineation of multiple-refracting
               horizons. Only selected raypaths for shotpoint 1, 2 and 3 are shown. The
               raypaths for shotpoints 4 and 5 are the mirror image (with respect to shotpoint 3)
               of the raypaths for shotpoints 2 and 1, respectively (from Haeni, 1988).
                                                                                           79.

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Formulas for solving problems involving more than two layers have been developed from this
relationship and other geometric properties (Mooney, 1984; Haeni, 1988; and Crice, 1992).
Refraction interpretation formula assumptions are:  1) layer boundaries are planar and have a
uniform slope; 2) the land  surface is flat; 3) all layers are seismically homogeneous and isotropic;
and 4) seismic velocity increases with depth. Formulas for three-layer problems can be readily
solved by hand or with hand-held calculators (Ballantyne and others, 1981).  More complex
formulas are solved using computers (Hunter, 1981; Mooney, 1984; and Orndorff and others,
1989).

Seismic refraction surveys are usually conducted by placing a cable and geophones along a linear
trend at regular spacings.  To maximize information, energy is sent into the ground from several
locations along the geophone spread and off-end. A typical layout of a seismic refraction line is
shown in Figure 7. The seismic wave sensed at each geophone is recorded on a seismograph
either as an analog or digital signal.  The array of geophones is physically or electronically moved
along the line to obtain additional seismic profiles. The spacing of the geophones is determined
in part by the depth of interest, the thickness of the layers and the size of the target.  There
should be at least two data points for each layer of interest to avoid ambiguities in interpretation.
Refraction cables commonly have takeouts for connecting the geophones at 25-, 50- and 100-foot
increments. Actual spacing can be adjusted by leaving slack in the cable between geophones.

For compressional wave seismic refraction surveys, the time of the arrival of the first compression
wave is selected for each geophone from the recording.  To enhance the signal from a low energy
source and to reduce the effect of noise, most modem seismographs allow stacking of repeated
shots.  Variations in the subsurface geology will cause the arrival times at the geophones to differ
and plots of the distance versus first-arrival time will be non-symmetrical.  Interpretations of the
layer dip, thickness, velocity and material types are based on these variations as seen in the time-
distance plots.

Interpretation of refraction data is done by plotting the first arrival times for each geophone on
the vertical axis at their distance from  the shot point (energy source) on the horizontal axis as
shown in Figure 8.  Straight lines are drawn through points thought to represent first arrivals at a
single layer.  The slope of a line is the inverse of the velocity of the layer it represents. Depth to
a layer interface is calculated based on either: 1) a crossover distance (XQ,, or x^) at which the
refracted wave traveling along the layer interface  reaches the geophone faster than the direct

80.

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wave; or 2) by projecting each segment to zero distance and using the zero intercept time (t2u, or
tM).  The crossover distance for each interface is found at the-break in slope as shown of Figure
8. The intercept times and the crossover distances are directly dependent on the thickness and
the compressional velocity of the each layer.  A detailed discussion of seismic refraction theory,
field methods,' and interpretation is given by Mooney (1984) and Haeni (1988).

An alternative method for interpreting seismic refraction data that give a depth and velocity
profile beneath each geophone is the generalized reciprocal method (GRM). GRM
interpretation calculates velocity analysis and time-depth functions to define  continuous forward
and reverse direction profiles using the technique of phantoming (Palmer, 1980 and 1981;
Lankston and Lankston, 1986; and Lankston, 1990). The computed values for the velocity
analysis and the time-depth functions are referenced to a position, termed the  G-position, which
lies half-way between two geophones separated by a distance XY.  Multiple calculations are done
for various XY distances to generate values of the velocity analysis and time-depth functions
versus G-position, creating a suite of curves.  Based on a subjective criteria of either minimal
curve irregularity for the velocity analysis function or maximum curve irregularity for time-depth
function, an optimum XY spacing is selected. The time-depth function curve for the optimum
XY spacing is then used to calculate the depth to the target refractor surface at each G-position.
The closer the geophones are spaced the more detailed the information on the refractor surface.

When an optimal XY distance can be determined with confidence, GRM can determine the
depth to the target refractor even when the velocity and thickness of the overlying layers are
unknown. GRM can also give an indication that undetected layers are present and indicate
whether they are hidden because of a velocity inversion or because they are too thin (Lankston,
1989 and 1990; and Palmer, 1980).  The GRM requires many calculations, and computer
programs are available (Hatherly, 1976).

Special refraction and downhole survey are performed for additional information about the
dynamic properties of the subsurface layers, including shear velocity and shear  modulus.  The
geophones and energy sources differ from normal compressional surveys and digital recording of
the wave form is required to process the signal (Mooney,  1974; and Dobecki,1979).
                                                                                      81.

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Shotpoint
                                                           9   N. 10     11     12
                                             DISTANCE
  Figure 8..     Typical geophone and shotpoint layout with seismic raypaths and time-distance
               plots for a two-layer model with a dipping boundary (from Haeni, 1988).
  82.

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Because the material properties measured by seismic methods, primarily velocities of sound
waves, are ,not directly related to AMD geochemistry or conditions that generate AMD, the
application of shallow seismic refraction to AMD investigations is limited  to general evaluations
of subsurface geology and hydrology. Knowledge about subsurface geology and hydrology can
complement and help calibrate other geophysical methods such as D;C. resistivity and
electromagnetics. Seismic refraction and perhaps reflection surveys can:

       o      provide complementary data for calibration of other geophysical surveys;
       o      delineate depth and lateral extent of waste fills;
       o      define and locate geologic structures such as faults and formation contacts;
       o      identify gravel channels within finer grained materials;
       o      define water table in unconsolidated sediments;
       o      identify subsurface voids; and
       o      azimuth refraction surveys can be used to measure anisotropy in fractured rock
              and in some cases determine the main direction of vertical fracturing.

Papers that describe the use of seismic refraction methods in environmental and hydrogeologic
investigations include: Burke (1967), Eaton and Watkins (1967), Lennox and Carlson (1967),
Bruehl (1983), Kopsick and Sanders (1983), Taylor and Cherkauer (1984), Underwood and others
(1984), Haeni (1988), King and others (1989), Barker (1990), Lankston (1990), Steeples and
Miller (1990), Carpenter and others (1991), Cooksley (1992), Crice (1992), and Davies and others
(1992).

Although reflection methods are generally not applied to shallow environmental or hydrogeologic
investigations, some research has been reported by Muir and Higgins (1990), Davies and King
(1992), Davies and others (1992), King (1992), Hill (1992), and Brabham and McDonald (1992).
                                                                                        83.

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                          GRAVITY GEOPHYSICAL METHODS14
          \
Measurement of the earth's gravitational field at the surface are compared to an arbitrary datum
to learn about subsurface structures. Predictable differences in the earth's gravitational field are
caused by changes in elevation and latitude, diurnal tidal distortion, and lateral variations in rock
bulk density.  The standard cgs unit for the earth's gravitational acceleration is approximately 980
cm/sec2 or 980 Gals (Galileos) = Ig (gravity), and 1 Gal  = 1 cm/sec2 = 0.0010197 g.  The units
of gravitational acceleration commonly used in land gravity surveys is milligals (mGals).  One
milligal is one-thousandth of a Gal = 1 x IQr6 g.  Gravity meters can detect variations in the
earth's gravitational field as small as a few hundredths of a milliGal or better.  For example, the
LaCoste  and Romberg Model D microgravity meters are precise to 0.001 mGal (1 x 10"9 g). In
order to  identify small geologic changes, field data must be corrected, to account for variations in
elevation, latitude, terrain, tidal effects and instrument drift (Telford and others, 1990).

Free-air and Bouguer elevation corrections must be applied to field data.  The free-air correction
adjusts for the difference in distance from the center of the earth to the datum plane and to the
station. The datum plane is usually sea level but adjustments  can be made to any datum.  The
adjustment is the addition of 0.09406 mGal per foot (03086 mGal/m) of elevation above the sea
level datum.  The adjustment is made regardless of whether or not there is rock material between
the sea level datum and the station, hence the term free-air correction.  The Bouguer correction
removes  the effect of the gravitational attraction caused by an assumed infinite slab of material
between  the horizontal plane of each station and the datum.  The Bouguer correction assumes an
infinite slab of material of bulk density pb with a base at the datum (sea level) and the top at the
station elevation. A Bouguer correction of 0.01278-pb mGal/ft (0.04192-pb mGal/m) is subtracted
from the field data, assuming a sea level datum (Telford and others, 1990).

Field data must be corrected for latitude because the earth is  not a true sphere, but an ellipsoid.
The latitude correction is subtracted from the free-air and Bouguer corrected data and is given
by the formula (Telford and others, 1990):

              g = 978031.846 • [1 +  0.005278895 •  sin2() + 0.000023462 • sin4(<|>)]
       ^Reference begin on page 133.
 84.

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where:                      g      =      acceleration of gravity, mGals
                            <(>      =      the latitude of station in degrees

The terrain correction adjusts for the effects of topographic variations in the vicinity of the data
station. Hills nearby the station give an upward component of gravitational attraction that
counteracts a portion of the downward pull of the earth. Valleys nearby have the effect of
causing a smaller downward pull at the station than is accounted for by the Bouguer correction.
The terrain correction is begun by centering a template over the station and  estimating the
average elevation within each area delineated by radial lines and concentric circles. The terrain
correction factor is the elevation difference between the station and each compartment, and is
calculating the gravitational effect represented by each compartment. The total terrain correction
factor is added to the gravity value at each station (Dobrin, 1976; and Telford and others, 1990).

If a high precision gravity survey is required, correction for the effects of tidal changes and
instrument drift must be made. Periodic readings at one station are necessary to estimate the
magnitude of the drift and adjust the data accordingly.  Earth's tide  can cause as much as 0.3
mGal cyclic changes in gravity at any one point (Telford and others, 1990).

The correction  of the field data for free-air, Bouguer, and latitude is termed  the Simple Bouguer
Anomaly.  The inclusion of the terrain correction is called the Complete Bouguer Anomaly.

One of the main problems with interpretation of gravity data is the need to remove the effects  of
other features, such as regional gradients, that are superimposed onto the signature of the
anomaly of interest.  Several methods that have been used  to extract the anomaly of interest
include the following methods (Telford and others, 1990):

o      The trend surface method assumes that regional gravitational effects can be modeled by a
       simple low-order polynomial  surface such as a plane. The trend surface is then
       subtracted from the observed data and a map of gravity residuals results.  The concept is
       that the removal of the regional data will enhance the trends caused by the local
       anomalies.  In some cases where data are noisy, a high-order polynomial surface is fit to
       the observed data before subtracting out the low-order regional surface.  Drawbacks to
       the method include that the regional effects may not be  accurately represented by a
       simple surface and that the use of high-order polynomials to fit the field data can
       introduce anomalies that are due solely to the fitting processes.
                                                                                       85.

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o      A spatial frequency filtering method uses band-pass filters to isolate spatial frequency
       components caused by shallow or deep density changes, This method requires some
       knowledge of the wavelength of the feature being investigated. The apparent wavelength
       of the anomaly is roughly proportional to the depth of the lateral density change causing
       the anomaly.  A shallow mass or near-surface noise causes a short wavelength anomaly
       while deep regional mass causes a broad wavelength anomaly.

o      The forward modeling method computes an ideal gravitational simulation that should be
       caused by the feature being investigated. Results of the simulated model are then
       compared with the actual data. Repeat simulations and comparisons are made until an
       acceptable fit is found.

o      The graphical method produces a smoothed surface that does not contain local anomalies
       and then this surface is subtracted from the original data to produce a map of residuals
       representing the local anomalies.

o      The gridding method is used to predict the regional gradient by averaging the Values
       within a specific radius of a station. The radius is on the same order of magnitude as the
       depth of the anomaly of interest.

o      The second derivative method uses a grid to calculate the second derivative at a station
       from a weighted average of surrounding stations. The second derivative method measures
       the curvature of the gravity field and is most influenced by shallow anomalies.

For environmental and engineering studies where the targets are generally small shallow features,

high-precision microgravity surveys are done.  Microgravity studies of small areas require precise

instrument readings and gravity station location. A precision of 10 centimeters in elevation and

30 meters in latitude (0.03 mGal accuracy) with grid spacings of a few meters  is commonly
needed (Telford and others, 1990).  Environmental surveys of small areas generally assume that
the regional gradient is a simple sloping plane or curve, like most geophysical methods,
interpretation of gravity data is non-unique and the interpreter must use all available data to limit

the solutions.
Application of gravity surveys in AMD investigations would include:  delineating waste pile

dimensions, especially for deep canyon fills; and identifying voids, tunnels and subsurface mine

workings. Reports and papers on the use of gravity surveys for environmental, engineering and

hydrogeologic studies include:  Dean (1958), Eaton and Watkins (1967); Lennox and Carlson

(1967); Arzi (1975), Ibrahim and Hinze (1972), Zohdy and others (1974), Carmichael and Henry

(1977), Butler (1984), Dahlstrand (1985), Roberts (1989), Roberts and others (1989), Hinze
(1988), Kick (1989), Adams and Hinze (1990), Hart and Muir (1990), Richard and Wolfe (1990),

Sandberg and Hall (1990),  Wolfe and Richard (1990), West (1992), and Kelly and Mares (1993).
 86.

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                         MAGNETIC GEOPHYSICAL METHODS15

Land based magnetic surveys measure the variations in the earth's magnetic field caused by near-
surface ferrous materials to locate buried targets.  The earth's total magnetic field intensity is not
uniform because of the dipoler nature of magnetism and because of the distribution in the earth
of rocks having  differing magnetic properties. The sun and solar flares temporarily distort the
earth's magnetic field.

Diurnal variations occur during daylight hours although their timing and magnitude are not
totally predictable (Telford and others, 1990). They usually must be removed from total field
intensity data because their magnitude of tens of gammas can obscure anomalies of interest.
During field surveys that last several hours,  a recording base station or periodic recording at the
same point is needed in order to record the diurnal variation.

Modern magnetic instruments usually measure the total magnetic field intensity. An alternative
method is the vertical gradient magnetometer survey. This method measures the magnetic field
at two pomts vertically separated by approximately 1 meter with the gradient being the change in
magnetism between sensors.  While the  total magnetic field for the earth varies throughout the
day, the gradient between these two vertically separated points should be relatively constant.
Thus there is no need to correct for diurnal variations, and repeat surveys should give similar
data.

The magnetometer survey is generally carried out along a linear traverse or a grid.  Data
corrected for diurnal effects and sometimes smoothed are generally analyzed qualitatively by
contouring or plotting a profile.  Quantitative interpretation can be done (Telford and others,
1990), but usually the buried metal objects of interest in environmental work do not lend
themselves to simple geometric models.

While there are several types of instruments used for measuring the earth's magnetic field, the
proton-precession magnetometer is the most common in environmental and engineering studies
(Breiner, 1973 and 1992; and Telford and others,  1990).  The proton-precession magnetometer
can read the total magnetic field to a sensitivity of 0.1 gamma.  The total intensity of the earth's
        References begin on page 135.
                                                                                      87.

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magnetic field in the contiguous United States ranges from 50,000 to 60,000 nanoteslas (1
nanotesla ^ 1 gamma = 10'5 gauss = 105 oersted = 10"9 web'ers/m2 = 10'9 tesla).

Magnetometer surveys can be of use in AMD investigations because they allow for rapid
identification o'f shallow buried  ferrous objects, common to mine sites.  Magnetometer surveys
can also aid in distinguishing between a conductive soil/rock mass from buried ferrous metal as
being the source of an electromagnetic, low resistivity anomaly.  For this reason, magnetometer
surveys are commonly done for quality control in electromagnetic and resistivity surveys.

Details on the theory and application of proton-precession magnetometers can be found in
Breiner (1973 and 1992) and Telford and others (1990), as well as general geophysical textbooks
(see list of general references).  Discussions of survey methods and interpretation for
environmental problems are given by Zohdy and others (1974), Fowler and Pasicznyk (1985),
Hinze (1988), Allen and Rogers (1989), and DeReamer and Pierce (1990).
                      GROUND PENETRATING RADAR METHOD16

In recent years the increased demand in the environmental studies for detailed knowledge about
the shallow subsurface and of objects buried within the upper 15 to 30 feet has led to the
development of a ground-based, high frequency, electromagnetic geophysical instrument known
as ground penetrating radar (GPR).  This geophysical instrument employs a short duration,
electromagnetic pulse using a broad-bandwidth antenna placed directly on the ground. The
depth of penetration and resolution are controlled by the antenna frequency which commonly
ranges from 80 Hz to 1000 Hz.  GPR detects differences in the dielectric properties of buried
materials. Voids and buried objects such as barrels and sewer lines have sufficiently different
dielectric properties to be detected.

GPR has advantages over other geophysical methods in that: 1) it can detect changes in  the
dielectric properties of material that may not be apparent to magnetic or resistivity geophysical
methods; 2) it has the highest target resolution of any shallow depth geophysical method; 3) a
survey can be done as rapidly as the operator can walk it; 4) output is a real-time plot and does
        6References begin on page 136.

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not require correction or processing; and 5) real-time graphic output allows rapid interpretation
and modification of survey.

GPR limitations include: 1) survey results are site dependent; 2) an increase in moisture content
of soils can reduce the signal, and ground water absorbs the signal; 3) an increase in clay content
reduces the signal; 4) penetration and resolution are frequency dependent; 5) the antenna must
pass near the object to detect it.

GPR surveys are typically conducted along a grid or a linear traverse across an area of interest.
Grid size varies with the target size (see previous section on line spacing and probability, page
24).  Preliminary surveys are needed to calibrate the instrument to site conditions.

GPR can be applied to AMD investigations where shallow buried objects are of interest and
contrasts in the dielectric properties of materials or voids are assumed.  These types of
investigations might include locating buried barrels, pipelines, old disposal trenches, and shallow
voids. A drawback to the use of GPR is the radar antenna must be in direct contact with the
ground to develop an adequate return signal. An irregular surface common at mine sites will
significantly affect the results.

Review  of GPR theory and methods for environmental studies is given by: Morley, (1974), Cook
(1975), Coon and others (1981), Benson and others (1984), Davis and others (1984), Wright and
others (1984), Underwood and Bales  (1984), Olhoeft (1988), Boucher and Galinovsky  (1989),
Davis and Annan (1989), Daniels (1989), Filler and Kuo (1989), Hennon (1990), Beres and
Haeni (1991), Fisher and others (1992), and Allen and Seelen (1992).
                         BOREHOLE GEOPHYSICAL METHODS17

Borehole geophysical or well logging includes all methods for lowering a sensing device down a
cased or uncased borehole to record physical, chemical, electrical, electromagnetic, or radioactive
parameters along the well bore.  Borehole logging is used extensively by the petroleum and
minerals industries, and to a limited extent by the geotechnical and water resources industries.  In
        References begin on page 138.
                                                                                       89.

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recent years with the increase in the number environmental and hydrogeologic investigations
associated with ground-water pollution, borehole geophysics have become an important tool to
assist the geologist and hydrogeologist in interpreting subsurface geology and hydrology.

Papers that report extensive use of borehole geophysical methods in the study of AMD problems
or in the monitoring of AMD were not found in the literature search.  Nevertheless, there are
numerous studies of high specific conductance ground-water pollution that utilized borehole
geophysics.  These studies clearly indicate that borehole geophysics can have direct application to
AMD investigations and to long-term monitoring of AMD pollution. There are numerous
borehole geophysical methods and many variations on each method developed for special
applications. The purpose of this summary is not to cover the topic in detail, but to provide a
general overview of the available methods and references for more detailed research. Therefore,
the following is a brief discussion of the methods of borehole geophysics as might be applied to
the investigation of AMD.

The types of borehole geophysical logging devices are numerous but can be generally divided into
the following five groups based on the parameter measured:

       o      Electrical resistivity and conductivity;
       o      Nuclear,
       o      Acoustic;
       o      Borehole  fluid; and
       o      Well construction.
Electrical resistivity or conductivity logging tools are designed to measure formation and
formation water resistance using a variety of designs and electrode configurations.  These
methods generally are done in an open hole although some of the electromagnetic induction
tools can penetrate non-conductive well casing, such as plastic.  The electrical methods are
accomplished by inducing current into the borehole wall and reading the response at one or more
electrodes.  The electrical resistivity methods include:  point resistivity, short normal, long
normal, lateral resistivity, focused resistivity, microlog, dipmeter, induced polarization,  and
electromagnetic induction logs. This group should also include a special type of electrical log,

90.

-------
 spontaneous potential, that measures natural potential differences caused by electrochemical
 reactions between layers.
           \

 Nuclear logging commonly uses measurements of either the level of natural radioactive material
 decaying in formations, or the level of backscatter or adsorption in the formation material caused
 by a radioactive source within the logging tool. The characteristics of the formations are
 determined by the amount of natural radioactivity or the response to radiation.  An advantage of
 radioactive logs is that they can be run in open or cased holes.  Commonly used radioactive logs
 include: natural gamma, gamma-gamma, neutron, spaced neutron, density, gamma spectrometry,
 neutron-activation, pulsed-neutron-decay, and dual-detector-density logs.

 Acoustic logging uses sound waves generated by the logging probe to measure formation
 properties, primarily porosity, fracture density and orientation, and the quality of annular space
 cementation. Acoustic methods include: sonic, tube wave amplitude, variable density, full
 waveform  sonic, borehole acoustic televiewer, and acoustic cement bond logs.

 Borehole fluid logs measure the physical and chemical characteristics of the fluids in the borehole
 and the adjacent formation and include: fluid conductivity, brine injector-detector, radioactive
 tracerjector, Eh, pH, temperature and differential-temperature, impeller-type flowmeter, and
 heat-pulse flowmeter logs. These logs must be run in an open hole that has ground water in
 equilibrium with the formation water.

 The final group are well-construction logs:  downhole video camera, casing-collar indicator,
 directional survey and caliper. These are run as part of the engineering design of the well or
 following well completion to document the construction.

 Interpretation of borehole logs can be done qualitatively or quantitatively. Most quantitative
 methods were developed by the petroleum and minerals industries. Numerous books are
 available on well log interpretation and include:  Lynch (1962), Schlumberger (1972a, 1972b,
 1974, 1977), Dresser Atlas (1975), Merkel (1979), Hearst and Nelson (1985), Labo (1986), and
 Brock (1986).

. Several papers that discuss the principles, applications and interpretation of borehole geophysics
 as it relates to environmental and fresh ground-water investigations are available and include:

                                                                                        91.

-------
Jones and Buford (1951), Keys (1967 and 1989), Keys and MacCary (1971), MacCary (1978),
Engineering Enterprises (1985), Driscoll (1987), Collier and Alger (1988), Stegner and Becker
(1988), Stowell (1989), Daniels and Keys (1990), Howard (1990b), Paillet and Saunders (1990),
Roscoe Moss Company (1990), Yearsley and Crowder (1990 and 1991), Crowder and others
(1991), Yearsley and others (1991), Yearsley and Crowder (1991), Mwenifumbo (1993), and
Welenco Inc. (undated).

Several authors address the use of borehole geophysical logs in determining subsurface
parameters of shallow formations and fresh-water aquifers. Parameters that are measured
include: lithology, porosity, permeability, bulk density, clay and shale content, moisture content,
ground water flow direction and velocity, and chemical dispersion.  Papers on these subjects
include: Croft (1971), Barker and Worthington (1973), Griffiths (1976),  Ogbe and Bassiouni
(1978), Biella and others (1983), Snelgrove and McNeill (1985), Spencer (1985), Huntley (1986),
Taylor and others (1989), Burns (1990), and Hess and Paillet (1990).

Correlation of borehole geophysical measurements and quality of shallow ground water is
discussed by numerous authors including: Turcan  (1962 and 1966), Alger (1966), Moore (1966),
Desai and Moore (1969), Worthington and Barker (1972), Worthington  (1976), Kwader (1985
 and 1986), Guo (1986), Alger and Harrison (1988), and Jorgensen (1989 and 1990).

 Studies on the use of borehole geophysical methods used to investigate fracture zones, fracture
 density, orientation, and transmissivity are given by:  Deluca and Buckley (1985), Morin and
 others (1988),  Merin (1989),  and Howard (1990a).

 Figure 9 is a matrix listing most of the geophysical logs commonly used in environmental and
 ground-water investigations along with the applicability of each method  to various geologic and
 hydrologic studies (Keys and MacCary, 1971; Welenco, undated).  Figure 9 indicates whether the
 logging method can be used with open or cased holes.

 Application of borehole geophysical logs to  the investigation of AMD is similar to the application
 to other environmental investigations.  Applications include: aid in identifying and correlating
 subsurface geologic units, determining physical parameters such as hydraulic conductivity,
. determining water quality, evaluating the interconnection of aquifers, and determining the
 effectiveness of ground-water cleanup programs.  An additional use of borehole logging for AMD

 92.

-------
investigations is the use of conductivity logs to map the extent and change in high specific
conductance ground-water pollution.  For example, electromagnetic induction tools, such as the
Geonic's EM39, can be run down wells with non-metal casings to measure the conductivity of the
formation fluids. Thus periodic logging of the well can monitor the appearance and extent of an
AMD plume.
                                                                                      93.

-------
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              engineering and environmental investigations.  Open circle = open hole; filled
              circle =  cased hole.  Casing material for EM methods must be non-conductive
              such as plastic,  (modified after Welenco, undated; and Keys and MacCary, 1971)
94.

-------
                                IV.  REFERENCES
                             GEOPHYSICAL TEXTBOOKS
Bhattacharyya, P. and Patra, H., 1968, Direct Current Geoelectrical Soundings: Elsevier Press,
    .   Amsterdam, the Netherlands, 135 p.

Dobrin, M.B., 1976, Introduction to Geophysical Prospecting (Third Edition):  McGraw-Hill
       Book Company, New York, New York, 630 p.

Griffiths, D.R, and King, R.F., 1981, Applied  Geophysics for Geologists and Engineers (Second
       Edition): Pergamon Press, New York, New York, 230 p.

Hearst, J.R., and Nelson, P.W., 1985, Well Logging of Physical Properties: McGraw-Hill, New
       York, New York, 571 p.

Keller, G.V., and Frischknecht, F.C., 1966, Electrical Methods in Geophysical Prospecting:
       Pergamon Press, New York, New York, 523  p.

Labo, J.,  1986, A practical  Introduction to Borehole Geophysics, Geophysical References,
       Volume 2: Society of Exploration Geophysicists, Tulsa, Oklahoma, 330 p.

LeRoy, L.W., and LeRoy, D.O., 1977, Subsurface Geology, Petroleum, Mining, Construction:
       Colorado School of Mines, Golden, Colorado, 941 p.

Lynch, EJ., 1962, Formation Evaluation: Harper and Row, Publishers, New York, New York, 422
       P-

Milsom, J., 1989, Field Geophysics-Geological  Society of London Handbook: Halsted Press, John
       Wiley and Sons, New York, New York, 182 p.

Nabighian, M. K, editor, 1988 (Volume 1-Theory), 1990 (Volume 2-Applications),
       Electromagnetic Methods in Applied Geophysics: Society of Exploration Geophysicists,
       Tulsa, Oklahoma.

Parasnis,  D.S., 1973, Mining Geophysics, 2nd Edition: Elsevier, New York, New York, 395 p.

Parasnis,  D.S., 1986, Principles of Applied Geophysics (Fourth Edition): Chapman Hall, New
       York, New York, 402 p.

Robinson, E.S., and Coruh, C, 1988, Basic Exploration Geophysics: John Wiley & Sons, New
       York, New York, 562 p.

Sharma, P.V., 1986, Geophysical Methods in Geology (Second Edition): Elsevier Science
       Publishing Co., New York,  New York,  442 p.
                                                                                   95.

-------
Sumner, J.S., 1976, Principles of Induced Polarization for Geophysical Exploration, Developments
       in Economic Geology, No. 5: Elsevier Science, Amsterdam, the Netherlands, 277 p.

Telford, W.M., Geldart, L.P. and Sheriff, R.E., 1976, Applied Geophysics: Cambridge University
       Press, New York, New York, 770 p.

Telford, W.M., Geldart, L.P. and Sheriff, R.E., 1990, Applied Geophysics (Second Edition):
       Cambridge University Press, New York, New York, 770 p.
 96.

-------
                                  REFERENCES  ON
                         ENVIRONMENTAL, GROUNDWATER
                    and GEOTECHNICAL GEOPHYSICAL METHODS


Association of Engineering Geologist, Sacramento Section, 1990, Geophysical Applications in
       Engineering Geology and Ground Water, 3rd Annual Seminar Proceeding, June 15-16,
       1990, Sacramento, California.

Benson, R.C., Turner, M, and Vogelson, W., 1988, In Situ, Time Series Measurements for Long
       Term Ground-Water Monitoring, in A.G. Collins and A.I. Johnson, editors, Ground-
       Water Contamination, Field Methods, ASTM STP-963: American Society for Testing and
       Materials, Philadelphia, Pennsylvania, p. 58-72.

Benson, R.C., Turner, M.S., Volgelson, W.D., and Turner, P.P., 1985, Correlation Between Field
       Geophysical Measurements and Laboratory Water Sample Analysis, in Proceedings of
       NWWA Conference on Surface and Borehole Geophysical Methods in Ground Water
       Investigation, February 12-14, 1985, Fort Worth, Texas: National Water Well
       Association, Worthington, Ohio, p. 178-197.

Boulding, J.R.,  1992, Use of Airborne, Surface and Borehole Geophysics at Contaminated Sites,
       EPA/625/R-92/007: U.S. Enviornmental Protection Agency, Office of Research and
       Development, Cincinnati.

Boulding, J.R.,  1993, Subsurface Characterization and Monitoring Techniques-A Desk Reference
       Guide, Volume 1, Solids and Ground Water, EPA/625/R-93/003A: U.S. Enviornmental
       Protection Agency, Office of Research and Development, Cincinnati.

Boulding, J.R.,  1993, Subsurface Characterization and Monitoring Techniques-A Desk Reference
       Guide, Volume 2, The Vadose Zone, EPA/625/R-93/003B: U.S. Enviornmental Protection
       Agency, Office of Research and Development, Cincinnati.

Bruehl, D.H., 1983, Use of Geophysical Techniques to Delineate Ground-Water Contamination,
       in Proceeding of the Third National Symposium on Aquifer Restoration and  Ground-
       Water Monitoring, May 25-26,1983, Fawcett Center, Columbus, Ohio: National Water
       Well Association, Worthington, Ohio, p. 295-300.

Glaccum, R.A., Benson, R.C., and Noel, M.R., 1982, Improving Accuracy and Cost Effectiveness
       of Hazardous Waste Site Investigations: Ground Water Monitoring Review, Summer
       1982, p. 36-40.

Greenhouse, J.P., Monier-Williams, M., 1985, Geophysical Monitoring of Ground Water
       Contamination Around Waste Disposal Sites: Ground Water Monitoring Review, Fall
       1985, p. 63-69.

Greenhouse, J.P., and Slaine, D.D., 1986, Geophysical Modelling and Mapping of Contaminated
       Groundwater Around Three Waste Disposal Sites in Ontario: Canadian Geotechnical
       Journal, v. 23, ho. 3, p. 372-384.
                                                                                  97.

-------
Griffiths, D.H., and King, R.F., 1981, Applied Geophysics for Geologists and Engineers (Second
       Edition): Pergamon Press, New York, New York, 230 .p.

Hitchcock, A.S., and Harmon, H.D., Jr., 1983, Application of Geophysical Techniques as a Site
       Screening Procedure at Hazardous Waste Sites, in Proceeding of the Third National
       Symposium on Aquifer Restoration and Ground-Water Monitoring, May 25-26,  1983,
       Fawcett Center, Columbus, Ohio: National Water Well Association, Worthington, Ohio,
       p. 307-312.

Kelly, W.E., Bogardi, I. Nicklin, M., and Bardossy, A., 1988, Combining Surface Geolelectrics and
       Geostatistics for Estimating the Degree and Extent of Ground-Water Pollution, in A.G.
       Collins and A.I. Johnson, editors, Ground-Water Contamination: Field Methods, ASTM
       STP-963: American Society for Testing and Materials, Philadelphia, Pennsylvania, p. 73-
       85.-

Kelly, W.E., and Mares, S., editors, 1993, Applied Geophysics in Hydrogeological and •
       Engineering Practice, in Developments in Water Science, v. 44: Elsevier Science
       Publishers B.V., New York, New York, 289 p.

Keys, W.S., 1989, Borehole Geophysics Applied to Ground Water Investigations: National Water
       Well Association, Dublin, Ohio, 313 p.

Keys, W.S., and MacCary, L.M., 1983, Application of Borehole Geophysics to Water Resource
       Investigations: Techniques in Water-Resource Investigations of the U.S. Geological
       Survey, Chapter El, Book 2,126 p.

Lewis, M.R., and Haeni, F.P., 1987, The Use of Surface Geophysical Techniques to  Detect
       Fractures in Bedrock-An Annotated Bibliography:  U.S. Geologic Survey Circular 987, 14
       P-

Lobo, J., 1986, A Practical Introduction to Borehole Geophysics, An Overview of Wireline Well
       Logging Principles for Geophysics, Geophysical References Series Volume 2: Society of
       Exploration Geophysicists, Tulsa, Oklahoma, 330 p.

Mazac, O., Kelly, W.E., and Landa, I., 1985, A Hydrogeophysical Model for Relations Between
       Electrical and Hydraulic Properties of Aquifers: Journal of Hydrology, v. 79, p.  1-19.

Merkel, R.H., 1986, Well Log Formation Evaluation, AAPG Continuing Education  Course Notes
       Series No. 14:  American Association of Petroleum Geologists, 82 p.

Mooney, H.M., 1980, Handbook of Engineering Geophysics, Volume 2, Electrical Resistivity:
       Bison Instruments, Minneapolis, Minnesota.

Mooney, H.M., 1984, Handbook of Engineering Geophysics, Volume 1, Seismic: Bison
       Instruments, Minneapolis, Minnesota.

Morley, L.W., 1967, editor, Mining and Groundwater Geophysics, Economic Geology Report  No.
       26, Geological'Survey of Canada, Proceedings of the Canadial Centennial Conference on
       Mining and Groundwater Geophysics, Niagara Falls, Canada, October 1967: Department
       of Energy, Mines and Resources, Ottawa, Canada, 722 p.
 98.

-------
 Orndorff, R.C., Dodd, K., Bunnells, G.B., and Sakss, Y, 1989, Computer Programs Released as
       U.S. Geological Survey Publications Through August 1989: U.S. Geological Survey Open-
       File Report 89-681, 70 p.

 Paillet, F.L., and Saunders, W.R., editors, Geophysical Applications for Geotechnical
       Investigations, ASTM STP-1101: American Society for Testing and Materials,
       Philadelphia,  Pennsylvania, 112 p.

 Saunders, W.R., and  Cox, S.A., 1988, Technical and Logistical Problems Associated with the
       Implementation and Integration of Surface Geophysical Methods in Inactive Hazardous
       Waste Site Investigations, in Proceedings of the Second National Outdoor Action
       Conference on Aquifer Restoration, Ground Water Monitoring and Geophysical
       Methods, Volume II, May 23-26,1988, Las Vegas, Nevada: Association of Ground Water
       Scientists,  Worthington, Ohio, p. 637-653.

 Stierman, D J., and Ruedisili, L.C., 1988, Integrating Geophysical and Hydrogeological Data:  An
       Efficient Approach to Remedial Investigation of Contaminated Ground Water, in A.G.
       Collins and A.I. Johnson, editors, Ground-Water Contamination: Field Methods, ASTM
       STP-963: American Society for Testing and Materials, Philadelphia, Pennsylvania, p. 43-
       57.

 Tweeton, D.R., Cumerlato, C.L., Hanson, J.C., and Kuhlman, H.L., 1991, Field Tests of
       Geophysical Techniques for Predicting and Monitoring Leach Solution Flow During In
       Situ Mining:  Geoexploration, v. 28, p. 251-268.

van Blaricom, R.,  compiler, 1992, Practical Geophysics II for the Exploration Geologist:
       Northwest Mining Association, Spokane Washington, 570 p.

Ward, S.H., editor, 1990, Geotechnical and Environmental Geophysics, in Investigations in
       Geophysics No. 5, Volume I, Review and Tutorial:  Society of Exploration Geophysicists,
       Tulsa, Oklahoma, 389 p.

Ward, S.H., editor, 1990, Geotechnical and Environmental Geophysics, at Investigations in
       Geophysics No. 5, Volume II, Environmental and Groundwater: Society of Exploration
       Geophysicists, Tulsa, Oklahoma, 343 p.

Ward, S.H., editor, 1990, Geotechnical and Environmental Geophysics, in Investigations in
       Geophysics No. 5, Volume III, Geotechnical: Society of Exploration Geophysicists, Tulsa,
       Oklahoma, 300 p.

Welenco, Inc., Water Well Geophysical Logs (Second Edition), Bakersfield, California, 81
       P-

Welenco, Inc., Water Well Geophysical Logs (Third Edition), Bakersfield, California, 59
       P-

Zohdy, A.A.R., Eaton, G.P., and Mabey, D.R., 1974, Applicaton of Surface Geophysics to
       Ground-Water Investigations: Techniques in Water-Resource Investigations of the U.S.
       Geological Survey, Chapter Dl, Book 2,116 p.
                                                                                     99.

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                                REFERENCES ON
             WATER QUALITY AND GEOPHYSiCAL METHODS
Agocs, W.B., 1955, Line Spacing Effects and Determination of Optimum Spacing Illustrated by
      Marmara, Ontario Magnetic Anomaly: Geophysics, v. XX, no. 4, p. 871-885.

Alger, R.P., 1966, Interpretation of Electric Logs in Fresh Water Wells in Unconsolidated
      Formations, in Transactions of the Society of Professional Well Log Analysts, Seventh
      Annual Logging Symposium, May 9-11, Tulsa, Oklahoma, p. CC1-CC23.

Barnes, I., and Clarke, F.E., 1964, Geochemistry of Ground Water in Mine Drainage Problems:
      U.S. Geological Survey Professional Paper 473-A, p. 6.

Barrow, G.M., 1979, Physical Chemistry, (Forth Edition): McGraw-Hill, New York, New York,

Benson, R.C., Turner, M., and Vogelson, W., 1988, In Situ, Time Series Measurements for Long
      Term Ground-Water Monitoring, in A.G. Collins and A.I. Johnson, editors, Ground-
      Water Contamination, Field Methods,  ASTM STP-963:  American Society for Testing and
      Materials, Philadelphia, Pennsylvania, p. 58-72.

Benson, R.C., Turner, M.S., Volgelson, W.D.,  and Turner, P.P., 1985, Correlation Between Field
      Geophysical Measurements and Laboratory Water Sample Analysis, in Proceedings of
      NWWA Conference on Surface and Borehole Geophysical Methods in Ground Water
      Investigation, February 12-14,1985, Fort Worth, Texas:  National Water Well
      Association, Worthington, Ohio, p.  178-197.

California State Water Resources Control Board, 1988, Mine Waste Study: University of
      California, Berkeley, July 1,1988, 416 p.

Carmichael, R.S., editor, 1989, Practical Handbook of Physical Properties of Rocks and Minerals:
      CRC Press, Boca Raton, Florida, 741 p.

Drever, J.I., 1982, The Geochemistry of Natural Waters (Second Edition): Prentice Hall,
      Englewood Cliffs, New Jersey, 437 p.

Frimpter, M.A., and  Maevsky, A., 1979, Geohydrologic Impacts of Coal Development in the
      Narragansett Basin, Massachusetts  and Road Island: U.S. Geological Survey Waster-
      Supply Paper 2062, 35 p.

Grady, S.J., and Haeni, F.P., 1984, Application of Electromagnetic Techniques in Determining
      Distribution and Extent of Ground Water Contamination at a Sanitary Landfill,
      Farmington, Connecticut, in D.M. Nelsen an M. Curl, editors, NWWA/EPA  Conference
      on Surface and Borehole Geophysical  Methods in Ground Water Investigations, February
      7-9,1984, San Antonio, Texas, p. 388-417.

Greenhouse, J.P., and Slaine, D.D., 1986, Geophysical Modelling and Mapping of Contaminated
      Groundwater Around Three Waste Disposal Sites in Southern Ontario: Canadian
      Geotechnical Journal, v, 23, p. 372-384.
100.

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Hem, J.D., 1989, Study and Interpretation of the Chemical Characteristics of Natural Water
       (Thjrd Edition): U.S. Geological Survey Water Supply-Paper 2254, 263 p.

Hill, H.J., and Milburn, J.D., 1956, Effects of Clay and Water Salinity on Electrochemical
       Behavior of Reservoir Rock, Transactions of the American Institute of Mining,
       Metallurgical and Petroleum Engineers (AIME), v. 207, p. 203.

Jorgensen, D.G., 1990, Estimating Water Quality From Geophysical Logs, in F.L. Paillet and
       W.R. Saunders, editors, Geophysical Applications for Geotechnical Investigations, STP
       1101: American Society for Testing and Materials, Philadelphia, Pennsylvania, p 47-64.

Keller, G.V., 1966, Electrical Properties of Rocks and Minerals, in Handbook of Physical
       Constants (Revised Edition): in S.P.  Clark, editor, Geological Society of America Memoir
       97, p. 553-557.

Keller, G.V., 1989, Electrical Properties, in R.S. Carmichael, editor, Practical Handbook of
       Physical Properties of Rocks and Minerals: CRC Press, Boca Raton, Florida, p. 361-427.

Keller, G.V., and Frischknecht, F.C., 1966, Electrical  Methods in Geophysical Prospecting:
       Pergamon Press, New York, New York, 515 p.

Kendal, M.G., and Moran, P.A.P., 1963, Geometrical Probability, in Griffins' Statistical
       Monographs and Courses, No. 10: Hafner Publishing Company, New York, New York,
       125 p.

Keys, W.S., 1989, Borehole Geophysics Applied to Ground Water Investigations: National Water
       Well Association, Dublin, Ohio, 313  p.

Keys, W.S., and MacCary, L.M., 1971, Application of Borehole Geophysics to Water-Resource
       Investigations: U.S. Geological Survey Techniques of Water-Resource Investigations,
       Book 2, Chapter El, 126 p.

Kwader, T., 1985, Resistivity-Porosity Cross  Plots for  Determining In-Situ Formation Water
       Quality-Case Examples, i/rProceedings of NWWA Conference on Surface and Borehole
       Geophysical Methods in Ground Water Investigation, February 12-14, 1985, Fort Worth,
       Texas: National Water Well Association, Worthington, Ohio, p. 415-424.

Laxen, D.P.H., 1977,  A Specific Conductance Method for Quality Control in Water Analysis:
       Water Research, v. 11, p. 91-94.

Lind, C.J., 1970, Specific Conductance as a Means of Estimating Ionic Strength: U.S. Geological
       Survey Professional Paper 700-D, p.  D272-D280.

Lynch, EJ., 1962, Formation Evaluation: Harper and Row Publishers, New York, New York, 422
       P-

Mayo, A.L., Nielsen,  P J., Loucks, M., and Brimhall, W.H., 1992, The Use of Solute and Isotopic
       Chemistry to Identify Flow Patterns  and Factors Which Limit Acid Mine Drainage in the
       Wasatch Range, Utah: Ground Water, v. 30, no. 2, p. 243-249.
                                                                                    101.

-------
Mazac, O., Kelley, W.E., and Landa, I., 1987, Surface Geoelectics for Groundwater Pollution and
       Protection Studies: Journal of Hydrology, v. 93, p. 277-294.

McNeill, J.D., 1980, Electrical Conductivity of Soils and Rocks, Technical Note TN-5: Geonics
       Limited, Ontario, Canada, 22 p.

Nordstrom, D.K., 1977, Hydrogeochemical And Microbiological Factors Affecting the Heavy
       Metal Chemisty of an Acid Mine Drainage System: PhD. Disseration, Stanford University,
       210 p.

Nordstrom, D.K., 1985, The Rate of Ferrous Iron Oxidation in a Stream Receiving Acid Mine
       Effluent, in Selected Papers in the Hydrologic Sciences 1985: U.S. Geological Survey
       Water-Supply Paper 2270, p. 113-119.

Onysko, S.J., 1985, Chemical Abatement of Acid Mine Drainage Formation: PhD. Disseration,
       University of California, Berkeley, 314 p.

Pankow, J.F., 1991, Aquatic Chemistry Concepts: Lewis Publishers, Chelsea, Michigan, 673 p.

Parasnis, D.S., 1986, Principles of Applied Geophysics (Fourth Edition): Chapman Hall, New
       York, New York, 402 p.

Rhoades, J.D., Raats, PAG, and Prather, RJ., 1976, Effects of Liquid-Phase Electrical
       Conductivity, Water Content, and Surface Conductivity on Bulk Soil Electrical
       Conductivity:  Soil Science Society of America Journal, v. 40, p. 651-655.

Sarma, V.VJ., and Roa, V.B., 1963a,  Variation of Electrical Resistivity of River Sands, Calcite
       and Quartz Powers with Water Content: Geophysics, v. 27, no. 4, p. 470-479.

Sarma, V.VJ., and Roa, V.B., 1963b, Reply to Discussion of their Paper "Variation of Electrical
       Resistivity of River Sands, Calcite and Quartz Powers with Water Content":  Geophysics,
       April 1963.

Singer, P.C., and Stumm, W., 1970, Acidic Mine Drainage: The Rate-Determining Step: Science,
       v. 167, February 20,1970, p. 1121-1123.

Slaine, D.D., and Greenhouse, J.P., 1982, Case Studies of Geophysical Contaminant  Mapping at
       Several Waste Disposal Sites,  in Proceeding of the Second National Symposium on
       Aquifer Restoration and Ground-Water Monitoring, May 26-28,  1982, Fawcett Center,
       Columbus, Ohio: National Water Well Association, Worthington, Ohio, p. 299-315.

Stumm, W., and Morgan, JJ.,  1981, Aquatic Chemistry  (Second Edition): John Wiley and Sons,
       New York, New York, 780 p.

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Worthington, P.P., 1975, Quantitative Geophysical Investigations of Granular Aquifers:
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                                                                                     117.

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       Ground Water Monitoring Review, Winter 1982, p. 46-51.

Lasky, L.R.,  1985, EM Conductivity for Leachage Plume Definition: A Case Study, in Abstracts
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Lawrence, T.A., and Boutwell, G.P., 1990, Predicting Stratigraphy at Landfill Sites Using
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       Theory and Practice, ASTM STP 1070: American Society for Testing and Materials,
       Philadelphia, Pennsylvania, p. 30-40.

Lyverse, M.A., 1989, Surface Geophysical Techniques and Test Drilling Used to Assess Ground-
       Water Contamination by Chlorides in an Alluvial Aquifer, in Proceedings on the Third
       National Outdoor Action Conference on Aquifer Restoration, Ground Water Monitoring
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       Florida: National Water Well Association, Dublin, Ohio, p. 993-1006.

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       Numbers, Technical Note TN-6:  Geonics Limited, Mississauga, Ontario,  Canada, 15  p.

McNeill, J.D., 1983, EM34-3 Survey Interpretation Techniques, Technical Note TN-8: Geonics
       Limited, Mississauga, Ontario, Canada, 7 p.

McNeill, J.D., 1985, EM34-3 Measurements at Two Inter-Coil Spacings to  Reduce Sensitivity to
       Near-Surface Material, Technical Note TN-19: Geonics Limited, Mississauga, Ontario,
       Canada, 3 p.
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McNeill, J.D., 1990, Use of Electromagnetic Methods for Groundwater Studies, in S.H. Ward,
       editor, Geotechnical and Environmental Geophysics, in Investigations in Geophysics No.
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       Oklahoma, p. 191-218.

McNeill, J.D., and Labson, V.,  1990, Geologic Mapping using VLF Radio Waves, in M. N.
       Nabighian, editor, Electromagnetic Methods in Applied Geophysics, Volume 2: Society of
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Mazac, O., Cislerova, M. Kelley, W.E., Landa, L, and Venhodova, D., 1990,  Determination of
       Hydraulic Conductivities by Surface Geoelectric Methods, in S.H. Ward, editor,
       Geotechnical and Environmental Geophysics, in Investigations in Geophysics No. 5,
       Volume II, Environmental and Groundwater: Society of Exploration  Geophysicists, Tulsa,
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Monier-Williams, M.E., Greenhouse, J.P., Mendes, J.M., and Ellert, N., 1990, Terrain
       Conductivity Mapping with Topographic Corrections at Three Waste Disposal Sites in
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       Investigations in Geophysics No. 5, Volume II, Environmental and Groundwater: Society
       of Exploration Geophysicists, Tulsa, Oklahoma, p. 41-55.

Morgenstem, KA., and Syverson, T.L., 1988, Determination of Contaminant Migration in
       Vertical Faults  and Basalt Rows with Electromagnetic Conductivity Techniques, in
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Parasnis, D.S.,  1973, Mining Geophysics (2nd Edition): Elsevier, New York, New York, 395 p.

Parasnis, D.S.,  1986, Principles of Applied Geophysics (Fourth Edition): Chapman Hall, New
       York, New York, 402 p.

Ruby, RJ., and Caoile, JA., 1984, Utilization of Shallow Geophysical Sensing at Two Abandoned
       Municipal/Industrial Waste Landfills on the Missouri River Floodplain: Ground Water
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Sartorelli, A.N., Pesowski, M.S., Wesselingh, L.G., 1990, Magnetic Induction Measurements in
       Support of Environmental Studies-Western Canada Examples: in Proceedings of 52nd
       EAEG Meeting and Technical Exhibition, May 28-June 1, 1990, Copenhagen.

Sanders, W.R., and Cox, S.A., 1988, Technical and Logistical Problems Associated with the
       Implementation and Integration of Surface Geophysical Methods in Inactive Hazardous
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Sharma, P.V., 1986, Geophysical Methods in Geology (Second Edition): Elsevier Science
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       Aquifer Restoration and Ground-Water Monitoring, May 26-28, 1982, Fawcett Center,
       Columbus, Ohio: National Water Well Association, Worthington, Ohio, p. 299-315.

Stewart, M. 1990, Rapid Reconnaissance  Mapping of Fresh-Water Lenses on Small Oceanic
       Islands, in S.H. Ward, editor, Geotechnical and Environmental Geophysics, in
       Investigations in Geophysics No. 5, Volume II, Environmental and Groundwater: Society
       of Exploration Geophysicists, Tulsa, Oklahoma, p. 57-66.

Stewart, M., and Bretnall, R., 1986, Interpretation of VLF Resistivity Data for Ground Water
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Street, G J., and Engel, R., 1990, Geophysical Surveys of Dryland Salinity, in S.H. Ward, editor,
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       Oklahoma, p. 187-199.

Syed, T., Zonge, K.L., Figgins, S., and Anzzolin, A.R., 1985, Application of Controlled Source
       Audio Magnetotellurics (CSAMT) Survey to Delineate Zones of Ground Water
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       Fort Worth,  Texas: National Water Well Association, Worthington, Ohio, p. 282-311.

Watzlaf, G.R., and Ladwig, KJ., 1987, Electromagnetic  Conductivity Surveys to Indentify Acid
       Sources and  Flow Patterns U.S. Coal Mines, in Proceedings, Acid Mine Drainage
       Seminar/Workshop Halifax, Nova Scotia, March 1987, p. 187-213.

Weber, D.D., Scholl, J.F., LaBrecque, D J., Walther, E.G., and Evans, R.B., 1984, Spactial
       Mapping of Conductive Ground Water Contamination with Electromagnetic Induction:
       Ground Water Monitoring Review, v. 4, no. 4, Fall 1984, p. 70-77.

Wightman, W.E., Martinek, B.C., and Hammermeister,  D., 1992, Geophysical Methods Used to
       Guide  Hydrogeologic Investigations at an UMTRA Site Near Grand Junction,  Colorado,
       in D.M. Nielsen and M.N. Sara, editors, Current Practices in Ground Water And Vadoze
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       Pennsylvania, p. 69-78.
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                                   REFERENCES ON
                      SELF POTENTIAL GEOPHYSICAL METHODS
Becker, A., and Telford, W.M., 1965, Spontanteous Polarization Studies: Geophysical
       Prospecting, v. 13, p. 173-188.

Bisdorf, R J., 1985, Electrical Techniques for Engineering Applications: Bulletin of Association of
       Engineering Geologists, v. XXII, no. 4, p. 421-433.

Bogoslovsky, V.A., Kuzmina, E.N., Ogilvy, A.A., and Strakhova, N.A., 1979, Geophysical
       Methods for Controlling The Seepage Regime in Earth Dams: Bulletin International
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Bogoslovsky, V.A., and Ogilvy A.A., 1970, Natural Potential Anomalies as a Quantitative Index
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Bogoslovsky, V.A., and Ogilvy A.A., 1972, The Study of Streaming Potentials on Fissured Media
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Bogoslovsky, V.A., and Ogilvy. A.A., 1973, Deformations of Natural Electric Fields Near
       Drainage Structures: Geophysical Prospecting, v.  321, no. 4, p. 716-723.

Bogoslovsky, VA., and Ogilvy A.A., 1977, Geophysical Methods for Investigation of Landslides:
       Geophysics, v. 42, no. 3, p. 562-571.

Butler, D.K., and  Llopis, J.L., 1990, Assessment of Anomalous Seepage Conditions, in S.H.
       Ward, editor, Geotechnical and Environmental Geophysics, in Investigations in
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Corwin, R.F., and Hoover, D.B., The Self-Potential Method in Geothermal Exploration:
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Erchul, R.A., 1988, The Use of Self Potential in the Detection of Subsurface Flow Patterns In
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Fitterman, D.V., 1976, Calculations of Self-Potential Anomalies Generated by Eh Potential
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Fitterman, D.V., 1979, Calculations of Self-Potential Anomalies Near Vertical Contacts:
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Fitterman, D.V., 1978, Electrokinetic and Magnetic Anomalies Associated with Dilatant Regions
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Furness, P., 1993, A Reconciliation of Mathematical Models for Spontaneous Mineralization
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Reznik, Y.M.,  1990,  Determination of Ground Water Paths Using Methods of Streaming
       Potentials, in 1990 National Symposium on Mining, University of Kentucky, Lexington,
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       Geophysics, v. 25, p. 226-249.

Satyanarayana Murty, B.V., and Haricharan, P., 1985, Nomogram for the Complete
       Interpretation of Spontaneous Potential Profiles Over Sheet-Like and Cylindrical Two-
       Dimensional Sources: Geophysics, v. 50, p. 1127-1135.

Schiavone, D., and Quarto, R., 1984, Self-Potential  Prospecting in the Study of Water
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Sill, W.R., 1983, Self-Potential Modelling From Primary Flows: Geophysics, v. 48, no. 1, p. 76-86.
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Stierman, D.J., 1984, Electrical Methods of Detecting Contaminated Groundwater at the
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                                                                                     129.

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                                  REFERENCES ON
                          SEISMIC GEOPHYSICAL METHODS
Ballantyne, EJ., Campbell, D.L., Menteraerier, S.H., and Wiggins, R., 1981, Manual of
       Geophysical Hand-Held Calculator Programs, Volume 2: Society of Exploration
       Geophysicists, Tulsa, Oklahoma.

Barker, R.D., 1990, Investigation of Ground Water Salinity by Geophysical Methods, in S.H.
       Ward, editor, Geotechnical and Environmental Geophysics, in Investigations in
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       Geophysicists, Tulsa, Oklahoma, p. 245-251.

Brabham, PJ.,-and McDonald, RJ., 1992, Imaging a Buried River Channel in an Intertidal Area
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Bruehl, D.H., 1983, Use of Geophysical Techniques to Delineate Ground-Water Contamination,
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       Surficial Deposits, in L.W. Morley, editor, Mining and Groundwater Geophysics,
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Cooksley, J.W., 1992, General Discussion of Seismic Methods, in R. van Blaricom, compiler
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Crice, D.B., 1992, Application for Shallow Exploration Seisographs, in R. van Blaricom, compiler,
       Practical Geophysics II for the Exploration Geologist: Northwest Mining Association,
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Davies, KJ., Barker, R.D., and King, R.F., 1992, Application of Shallow Reflection Techniques
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Davies, KJ., and King, R.F.,  1992, The Essentials of Shallow Reflection Data Processing:
       Quarterly Journal of Engineering Geology, v. 25, p. 191-206.

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      of Energy,  Mines and Resources, Ottawa, Canada, 544-568.

Haeni, P.P., 1988,  Applications of Seismic-Refraction Techniques to Hydrologic Studies:
      Techniques of Water-Resource Investigations of the U.S. Geological Survey, Chapter D2,
      Book 2, 86 p.

Hatherly, PJ., 1976, A Fortran IV Programme for the Reduction and Plotting of Seismic
      Refraction Data Using the Generalized Reciprocal Method: Report of the Geologic
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Hill, LA., 1992, Field Techniques and Instrumentation in Shallow Seismic Reflection: Quarterly
      Journal of Engineering Geology, v. 25, p. 183-190.

Hunter,  J.H., 1981, Software Listing of Programs for Shallow Seismic Exploration Using Apple
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King, R.F., 1992, High-Resolution Shallow Seismology: History, Principles and Problems:
      Quarterly Journal of Engineering Geology, v. 25, p. 177-182.

King, W.C., Witten, AJ., and Read, G.D., 1989, Detection and Imaging of Buried Wastes Using
      Seismic Propagation, Journal of Environmental Engineering: Americal Society of Civil
      Engineers, v. 115, p. 527-540.

Kopsick, D.A., and Sander, T.W., 1983, Refinement of the Shallow Seismic Reflection Technique
      in Determining Subsurface Alluvial Stratigrtaphy, in Proceeding of the Third National
      Symposium on Aquifer Restoration and Ground-Water Monitoring, May 25-26,1983,
      Fawcett  Center, Columbus, Ohio: National Water Well Association, Worthington, Ohio,
      p. 301-306.

Lankston, R.W., 1989, The Seismic Refraction Method: A Viable Tool for Mapping Shallow
      Targets Into the 1990s: Geophysics, v. 54, p. 1535-1542.

Lankston, R.W., 1990, High-Resolution Refraction Seismic Data Acquisition and Interpretation,
      in S.H. Ward, editor, Geotechnical and Environmental Geophysics, in Investigations in
      Geophysics No. 5, Volume I, Review and Tutorial: Society of Exploration Geophysicists,
      Tulsa, Oklahoma, p. 45-73.

Lankston, R.W., and Lankston, M.M., 1986, Obtaining Multilayer  Reciprocal Times Through
      Phantoming: Geophysics, v. 51, p. 45-49.

Lennox, D.H., and Carlson, V., 1967, Integration of Geophysical Methods for Groundwater
      Exploration in the Prairie Provinces, Canada, in Morley, L.W.,  editor, Mining and
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        File Report 89-681, 70 p.

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        Interpretation: Geophysics, v. 46, p. 1508-1518.

' Palmer, D., 1991, The Resolution of Narrow Low-Velocity Zones with  the Generalized
        Reciprocal Method: Geophysical Prospecting, v. 39, p. 1031-1060.

 Steeples, D.W., and Miller, R.D., 1990, Seismic Reflection Methods Applied to Engineering,
        Environmental, and Groundwater Problems, in S.H. Ward, editor, Geotechnical and
        Environmental Geophysics, in Investigations in Geophysics No.  5, Volume I, Review and
        Tutorial: Society of Exploration Geophysicists, Tulsa, Oklahoma, p.  1-30.

 Underwood, J.E., Laudon KJ., and Laudon, T.S., 1984, Seismic and Resistivity Investigations
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        91.

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                                   REFERENCES ON
                          GRAVITY GEOPHYSICAL METHODS
Adams, J.M., and Hipze, WJ., 1990, The Gravity-Geologic Technique of Mapping Buried
      Bedrock Topography, in S.H. Ward, editor, Geotechnical and Environmental Geophysics,
      in Investigations in Geophysics No. 5, Volume III, Geotechnical: Society of Exploration
      Geophysicists, Tulsa, Oklahoma, p. 99-105.

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

Butler, D.K., 1984, Microgravimetic and Gravity Gradient Techniques for Detection of
      Subsurface Cavities: Geophysics, v. 49, p. 1084-1096.

Carmichael, R.S., and Henry,  G., 1977, Gravity Exploration for Groundwater and Bedrock
      Topography in Glaciated Areas: Geophysics, v. 42, p. 850-859.

Dahlstrand, T.K., 1985, Applications of Microgravity Surveys to Subsurface Exploration, in
      Proceedings of NWWA Conference on Surface and Borehole Geophysical Methods in
      Ground Water Investigation, February 12-14, 1985, Fort Worth, Texas: National Water
      Well Association, Worthington, Ohio, p. 85-101.

Dean, W.C., 1958, Frequency Analysis for Gravity and Magnetic Interpretations: Geophysics, v.
      23, p. 97-127.

Dobrin, M.B., 1976,  Introduction to Geophysical Prospecting (Third Edition): McGraw-Hill Book
      Company, New York,  New York, 630 p.

Eaton, G.P., and Watkins, J.S., 1967, The Use of Seismic Refraction and Gravity Methods in
      Hydrogeological Investigation, in L.W. Morley, editor, Mining and Groundwater
      Geophysics, Economic Geology Report No. 26, Geological Survey of Canada: Department
      of Energy, Mines and  Resources, Ottawa, Canada, p. 544-568.

Hart, W.S., and Muir, S.G., 1990, Gravity as a Tool for Engineering and Environmental Geology,
      Applied Studies, in Geophysical Applications in Engineering Geology and Ground Water,
      3rd Annual Seminar Proceeding, June 15-16,1990, Sacramento, California: Association of
      Engineering  Geologist, Sacramento Section.

Hinze, WJ., 1988, Gravity and Magnetic Methods Applied to Engineering and Environmental
      Problems, in Proceedings of the Symposium on the Application of Geophysics to
      Engineering  and Environmental Problems (SAGEEP) March 13-16, 1989: Society  of
      Engineering  and Mineral Explorationm Geophysicists, Golden, Colorado, p. 1-107.

Ibrahim, A., and Hinze, WJ., 1972, Mapping Buried Bedrock Topography with Gravity: Ground
      Water, v. 10, p. 18-23.

Kelly, W.E., and Mares, S., editors, 1993, Applied Geophysics in Hydrogeological and
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       Publishers B.V., New York, New York, 289 p.
                                                                                   133.

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Kick, J.F., 1989, Landfill Investigations in New England Using Gravity Methods, in Proceedings
       of tfie Symposium on the Application of Geophysics to'' Engineering and Environmental
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       Geophysicists, Golden, Colorado, p. 339-353.

Lennox, D.H., and Carlson, V., 1967, Integration of Geophysical Methods for Groundwater
       Exploration in the Prairie Provinces, Canada, in L.W. Morley, editor, Mining and
       Groundwater Geophysics, Economic Geology Report No. 26, Geological Survey of
    •   Canada: Department of Energy, Mines and Resources, Ottawa, Canada, p. 517-535.  .

Richard, B.H., and Wolfe, P.H., 1990, Gravity as a Tool to Delineate Buried Valleys, in
       Proceedings of the Symposium on the Application of Geophysics to Engineering and
       Environmental Problems (SAGEEP) March 13-16,1989: Society of Engineering and
       Mineral Explorationm Geophysicists, Golden, Colorado, p. 59-105.

Roberts, R.L., 1989, A Multi-Technique Geophysical Approach to the Study of Landfills and
       Potential Ground Water Contamination, Masters Thesis: Purdue University, 247 p.

Roberts, R.L., Hinze, W J., and Leap, D.I., 1989, A Multi-Technique Geophysical Approach to
       Landfill Investigations, in Proceedings on the Third National Outdoor Action Conference
       on Aquifer Restoration, Ground Water Monitoring and Geophysical  Methods, May 22-25,
       1989, Orange County Convention Center, Orlando, Florida: National Water Well
       Association, Dublin, Ohio, p. 797-811.

Sandberg, S.K., and Hall, D.W., 1990, Geophysical Investigation of an Unconsolidated  Coastal
       Plain Aquifer System and the Underlying Bedrock Geology in Central New Jersey, in S.H.
       Ward, editor, Geotechnical and Environmental Geophysics, in Investigations in
       Geophysics No. 5, Volume II, Environmental and Groundwater: Society of Exploration
       Geophysicists, Tulsa, Oklahoma, p. 311-320.

Telford, W.M., Geldart, L.P. and Sheriff, R.E., 1990, Applied Geophysics (Second Edition):
       Cambridge University Press, New York, New York, 770 p.

West, R.E., 1992, The Land Gravity Exploration Method, in  R. van Blaricom, Practical
       Geophysics II for the Exploration Geologist, compiler: Northwest Mining Association,
       Spokane Washington, p. 177-233.

Wolfe, P J., and Richard, B.H., 1990, Geophysical Studies of Cedar Bogs, in  S.H. Ward, editor,
       Geotechnical and Environmental Geophysics, in Investigations in Geophysics No. 5,
       Volume II, Environmental and Groundwater: Society of Exploration  Geophysicists, Tulsa,
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134.

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                                  REFERENCES ON
                         MAGNETIC GEOPHYSICAL METHODS
Allen, R.P., and Rogers, B.A., 1989, Geophysical Surveys in Support of a Remedial
       Investigation/Feasibility Study at the Municipal Landfill in Metamora, Michigan, in
       Proceedings on the Third National Outdoor Action Conference on Aquifer Restoration,
       Ground Water Monitoring and Geophysical Methods, May 22-25,  1989, Orange County
       Convention Center, Orlando, Florida: National Water Well Association, Dublin, Ohio, p.
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Breiner, S., 1973, Applications Manual for Portable Magnetometers: EG&G Geometries,
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Breiner, S., 1992, Applications for Portable Magnetometers, in R. van Blaricom, compiler,
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DeReamer, J., and Pierce, D., 1990, Geophysical Investigation for Buried  Drums: A Case Study,
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       Seminar Proceeding, June 15-16,1990, Sacramento, California: Association of
       Engineering Geologist, Sacramento Section.

Fowler, J.W., and Pasicznyk, D.L., 1985, Magnetic Survey Methods Used in the Initial
       Assessment of a Waste Disposal Site, in Proceedings of NWWA Conference on Surface
       and Borehole Geophysical Methods in Ground Water Investigation, February 12-14,1985,
       Fort Worth, Texas: National Water Well Association, Worthington, Ohio, p. 267-280.

Hinze, WJ., 1988, Gravity and  Magnetic Methods Applied to Engineering and Environmental
       Problems, in Proceedings of the Symposium on the Application of Geophysics to
       Engineering and Environmental Problems (SAGEEP) March 13-16,1989: Society of
       Engineering and Mineral Explorationm Geophysicists, Golden, Colorado, p. 1-107.

Telford, W.M., Geldart, L.P. and Sheriff,  R.E., 1990, Applied Geophysics  (Second Edition):
       Cambridge University Press, New  York, New York, 770 p.

Zohdy, A.A.R., Eaton, G.P., and Mabey,  D.R., 1974, Application of Surface Geophysics to
       Ground-Water Investigations: Techniques of Water-Resource Investigations of the U.S.
       Geological Survey, Book 2, Chapter Dl, 116 p.
                                                                                  135.

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                                  REFERENCES ON
              GROUND PENETRATING RADAR GEOPHYCICAL METHOD
         *

Allen, R.P., and Seelen, M.A., 1992, The Use of Geophysics in the Detection of Buried Toxic
       Agents at a U.S. Military Installation, in D.M. Nielsen and M.N. Sara, editors, Current
       Practices in Ground Water And Vadoze Zone Investigations, ASTM STP-1118: American
       Society for Testing and Materials, Philadelphia, Pennsylvania, p. 59-68.

Benson, R.C., Glaccum, R.A., and Noel, M.R., 1984, Geophysical Techniques for Sensing Buried
       Wastes and Waste Migration, USEPA Contract No. 68-03-3053: Environmental
       Monitoring Systems Laboratory, Office of Research and Development, U.S.
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Beres, Jr., M., and Haeni, F.P., 1991, Application of Ground-Penetrating Radar Methods in
       Hydrogeologic Studies: Ground Water, v. 29, p. 375-386.

Boucher, R., and Galinovsky, L., 1989, RADAN 3.0, Geophysical Survey Systems, Inc., North
       Salem, New Hampshire.

Cook, J.C., 1975, Radar Transparencies of Mine and Tunnel Rocks: Geophysics, v. 40, p, 865-885.

Coon, J.B., Fowler, J.C., and  Schafers, CJ., 1981, Experimental Uses of Short-Pulse Radar in
       Coal Seams: Geophysics, v. 46, p. 1163-1168.

Daniels, J.D., 1989, Fundamental of Ground Penetrating Radar, in Proceedings of the
       Symposium on the Application of Geophysics to Engineering and Environmental
       Problems (SAGEEP) March 13-16,1989: Society of Engineering and Mineral Exploration
       Geophysicists, Golden, Colorado, p. 62-142.

Davis, J.L., and Annan, A.P., 1989, Ground-Penetrating Radar for High-Resolution Mapping of
       Soil and Rock Stratigraphy: Geophysical Prospecting, v. 37, p.  531-551.

Davis, J.L., Killey, R.W.D., Annan, A.P., and Vaughan, CJ., 1984, Surface and Borehole Ground-
       Penetrating Radar Surveys for Mapping Geologic Structure, in Proceeding of NWWA
       Conference on Surface and Borehole Geophysical Methods in Ground Water
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Filler, D.M., and Kuo, S.S., 1989, Subsurface Cavity Exploration Using Non-Distructive
       Geophysical Methods, in Proceedings on the Third National Outdoor Action Conference
       on Aquifer Restoration, Ground Water Monitoring and Geophysical Methods, May 22-25,
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       Association, Dublin, Ohio, p.  827-840.

Fisher, E., McMechan, G.A., and Annan, A.P., 1992, Aquisition and Processing of Wide-
       Aperature Ground-Penetrating Radar Data: Geophysics, v. 57, p. 495-504.

Hennon, K. 1990, Zillion Uses of Ground Penetrating Radar, in Geophysical Applications in
       Engineering Geology and Ground Water, 3rd Annual Seminar Proceeding, June 15-16,
       1990, Sacramento, California: Association of Engineering Geologist, Sacramento Section.
136.

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Morey, R.M., 1974, Continous Subsurface Profiling by Impulse Radar, in Proceeding of
       Engineering Foundations Conference on Subsurface Explorations for Underground
       Excavations and Heavy Construction, Henniker, New Hampshire, p. 213-232.

Olhoeft, G.R., 1988, Selected Bibliography on Ground Penetrating Radar, in Proceeding on
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       520.

Underwood, I.E., and Eales, J.W., 1984, Detecting a Buried Crystalline Waste Mass with Ground
       Penetrating Radar, in D.M. Nielsen and M. Curl, editors, NWWA/EPA Conference on
       Surface and Borehole Geophysical Methods in Ground Water Investigations: National
       Water Well Association, Worthington, Ohio, p. 654-665.

Wright, D.L., Olhoeft, G.R., and Watts, R.D., 1984, Ground Penetrating Radar Studies in Cape
       Cod, in D.M. Nielsen and M. Curl, editors, NWWA/EPA Conference on Surface and
       Borehole Geophysical Methods in Ground Water Investigations: National Water Well
       Association, Worthington, Ohio, p. 368-382.
                                                                                    137.

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                                  REFERENCES ON
                        BOREHOLE GEOPHYSICAL METHODS
Alger, R.P., 1966, Interpretation of Electric Logs in Fresh Water Wells in Unconsolidated
       Formations, in Transactions of the Society of Professional Well Log Analysts, Seventh
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Algers, R., and Harrison, C, 1988, Improved Fresh Water Assessment in Sand Aquifers, in
       Proceedings of the Second National Outdoor Action Conference on Aquifer Restoration,
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       Nevada: National Water Well Association,  Dublin, Ohio, p. 939-967.

Barker, R.D., and Worthington, P.E., 1973, Some Hydrogeophysical Properties of the Bunter
       Sandstone of Northwest England: Geoexploration, v. 11, p. 151-170.

Biella,  G., Lozej, A., and Tabacco, I., 1983, Experimental Study of Some Hydrogeophysical
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Brock,  J., 1986, Applied Open-Hole Log Analysis,  in Contributions in Petroleum Geology and
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Bums,  D.R., 1990, Acoustic Waveform Logs and the In-Situ Measurement of Permeablity, in F.L.
       Paillet and W.R. Saunders, editors, Geophysical Applications for Geotechnical
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Collier, H.A., and Alger, R.P., 1988, Recommendation for Obtaining Valid Data From Borehole
       Geophysical Logs, in Proceedings of the Second National Outdoor Action Conference on
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Croft, M.G., 1971, A Method of Calculating Permeability from Electric Logs, in U.S. Geological
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Crowder, R.E., Lo Coco, JJ., and Yearsley, E.N.,  1991, Application of Full Waveform Borehole
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Daniels, JJ., and Keys, W.S., 1990, Geophysical Well Logging for Evaluating Hazardous Waste
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Deluca, RJ., and Buckley, B.K., 1985, Borehole Logging to Delineate  Fractures in a
       Contaminated Bedrock Aquifer, in  Proceedings of NWWA Conference on  Surface and
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Desai, K.P., and Moore, E.J., 1969, Equivalent NaCl Determination from Ionic Concentrations:
       The Log Analyst, May-June, 1969, p. 12-21.

Dresser Atlas, 1975, Log Interpretation Fundamentals, Houston, Texas.

Driscoll, F.G., 1987, Ground Water and Wells (2nd Edition): Johnson Division, Saint Paul,
       Minnesota, 1089 p.

Engineering Enterprises, Inc., 1985, Log Interpretation Workshop for U.S. EPA Region IX,
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Griffiths, D.,  1976, Application of Electrical Resistivity Measurements for the Determination of
       Porosity and Permeabilty in Sandstones: Geoexploration, v. 14,  p. 207-219.

Guo, Y.A., 1986, Estimation of TDS in Sand Aquifer Water Through Resistiviy Log: Ground
       Water, v. 24, no. 5, p. 598-600.

Hearst,  J.R., and Nelson, P.W., 1985, Well Logging of Physical Properties: McGraw-Hill, New
       York, New York.

Hess, A.E., Paillet, F.L., 1990, Application of the Thermal-Pulse Flowmeter in the Hydraulic
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Howard, K.W.F., 1990a, Geophysical Well Logging Methods for Detection and Characterization
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Howard, K.W.F., 1990b, The Role of Well Logging in Contaminant Transport Studies, in S.H.
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Huntley, D., 1986, Relations Between Permeability and Electrical Resistivity in Granular
       Aquifers: Ground Water, v. 24, no. 4, p. 466-474.

Jones, P.H., and Buford, T.B., 1951, Electrical Logging Applied to Ground Water Exploration:
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Jorgensen, D.G., 1989, Using Geophysical Logs to Estimate Porosity, Water Resistivity, and
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Jorgensen, D.G., 1990, Estimating Water Quality From Geophysical Logs, in F.L. Paillet and
       W.R.  Saunders, editors, Geophysical Applications for Geotechnical Investigations, STP-
       1101:  American Society for Testing and Materials, Philadelphia, Pennsylvania, p. 47-64.
                                                                                    139.

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Keys, W.W., 1967, Borehole Geophysics as Applied to Groundwater, in L.W. Morley, editor,
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       614.

Keys, W.S., 1989, Borehole Geophysics Applied to Ground Water Investigations: National Water
       Well Association, Dublin, Ohio, 313 p.

Keys, W.S., and MacCary, L.M., 1971, Application of Borehole Geophysics to Water-Resource
       Investigations: U.S. Geological Survey Techniques of Water-Resource Investigations,
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Kwader, T., 1985, Resistivity-Porosity Cross Plots for Determining In-Situ Formation Water
       Quality-Case Examples, in Proceedings of NWWA Conference on Surface and Borehole
       Gepphysical Methods in Ground Water Investigation, February 12-14, 1985, Fort Worth,
       Texas: National Water Well Association, Worthington, Ohio, p. 415-424.

Kwader, T, 1986, The Use of Geophysical Logs for Determining Formation Water Quality:
       Ground Water, v. 24, no. 1, p. 11-15.

Labo, J., 1986, A practical Introduction to Borehole Geophysics, Geophysical References Volume
       2: Society of Exploration Geophysicists, Tulsa, Oklahoma, 330 p.

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Merin,  I.S., 1989, Characterization of Fractures in Devonian Siltsones, Northern Appalachian
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       Outdoor Action Conference on Aquifer Restoration, Ground Water Monitoring and
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Merkel, R.H., 1979, Well Log Formation  Evaluation, Continuing Education  Course Notes, Series
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Morin, R.H., Hess, A.E., and Paillet, F.L., 1988, Determing the Distribution of Hydraulic
       Conductivity in a Fractured Limestone Aquifer by Simultaneous Injection and
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Mwenifumbo, CJ., 1993, Borehole Geophysics  in Environmental Mining: Canadian Institute of
       Mining Bulletin, v. 86, no. 966, January 1993, p. 43-49.
 140.

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 Ogbe, D., and Bassiouni, Z., 1978, Estimation of Aquifer Permeabilities From Electric Logs: The
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       National Water Well Association, Dublin,  Ohio, p. 891-907.

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                                                                                   141.

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Turcan, A.N., 1966, Calculation of Water Quality from Electrical Logs, Theory and Practice,
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Worthington, P.P., and Barker, R.D., 1972, Methods for the Calculation of True Formation
       Factors in the Bunter Sandstone of Northwest England: Engineering Geology, v. 6, p.
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Yearsley, E.N., and Crowder, R.F., 1990, State-of-the-Art Borehole Geophysics Applied to
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Yearsley, E.N., and Crowder, R.F., 1991, State-of-the-Art Borehole Geophysics Applied to
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       Investigations, June 17-18,1991, Emeryville, California: COLOG Inc., Golden, Colorado,
       17 p.

Yearsley, E.N., Crowder, R.F., and Irons, L.A., 1991, Monitoring Well Completion Evaluation
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       1, Winter 1991, p. 103-111.
 142.

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V.    LIST OF ABBREVIATIONS
  A
  B
  °C
  cgs
  cm
  E
  3
  H
  Hz
  °K
  kJ
  L
  m
  mg
  mg/L
  raks
  mho
  mraho/m
  mm
  mol.
  mS
  mS/m
  mV
  Pa
  s
  S
  T
  v
  P.
  a
  «t>
  a
  AtS/cm
  /Ltmho/cm
  /xmho/m
ampere
magnetic field
degrees Celsius
centimeters-grams-seconds
centimeter
electrical field strength
Faraday's constant
magnetic field strength
Hertz = cycles / second
degrees Kelvin
kilojoules
liter
meter
milligram
milligrams / Liter
meters-kilograms-seconds
I/ohms  =   1/ii
millimhos / meter
millimeter
mole
milliSiemen
milliSiemens / meter
milliVolt
Pascal
second
longitudinal conductance
transverse resistance
volts
resistivity
conductivity
phase difference between E and H
ohm
magnetic permeability
magnetic field
microSiemen
microSiemens / centimeter
micromhos / centimeter
micromhos / meter
                                                          143.

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