Magnetic Methods

Magnetism has been studied for a very long time in human history.  Early Greek philosophers new about the attraction of iron to a magnet.  The first magnets consisted of a naturally occurring rock called lodestone, a variety of massive magnetite (almost pure iron oxide).  Magnetite is the only naturally occurring mineral with distinctly obvious magnetic properties.  Only a few other minerals have any detectable magnetism.  However, extremely sensitive magnetometers can detect trace magnetism in many different minerals.  Iron, because of its atomic structure, has the greatest tendency to become magnetized or aligned.  Other elements, such as cobalt and nickel, also have lesser tendency to become magnetic.  Any mineral or rock which contains any of these elements is likely be more magnetic.

Magnetism occurs when like poles of adjacent individual atoms are in alignment, creating a “dipole” effect.  Another word for this alignment effect is called “polarization”.  An analogy can be made with the north and south poles of a typical bar magnet.  Poles having the same charge repel each other; poles oppositely charged attract each other.  At extreme temperatures, the vibrations of the atoms cause a loss of alignment, leading to a loss of magnetism.   Experiments have shown this loss of magnetism occurs at a temperature of approximately 550^o C (also known as the “Curie point”).

Magnetic Minerals

Magnetic strength of a mineral or rock is therefore a function of two things:  1) the amount of iron, nickel or cobalt, and 2) the amount of alignment which takes place.  The measure of magnetic strength of a mineral or rock is called the “magnetic susceptibility” (Table 14 - 1).  This can be measured qualitatively with a simple magnet by testing the “pull”.  It can also be measured with a “magnetometer”.  The susceptibility of a completely nonmagnetic substance is equal to 0.  The susceptibility of a highly magnetic mineral (such as magnetite) is about 20.  Every mineral has at least some very minor amount of magnetic susceptibility, but for most minerals it is virtually nil.

Rock/Mineral	Magnetic Susceptibility
Rocks
Salt	0 – 0.001
Slate	0 – 0.002
Limestone	0.00001 – 0.0001
Granulite	0.0001 – 0.05
Rhyolite	0.00025 – 0.001
Greenstone	0.0005 – 0.001
Basalt	0.001 – 0.1
Gabbro	0.001 – 0.1
Dolerite	0.01 – 0.15
Minerals
Pyrite	0.0001 – 0.005
Hematite	0.001 – 0.0001
Pyrrhotite	0.001 – 1.0
Chromite	0.0075 – 1.5
Magnetite	0.1 – 20.0

The magnetic instrument we are most familiar with is the compass.  A piece of lodestone suspended from a string will align itself with the earth’s magnetic field.  In the 12^th century the Chinese discovered that rubbing a needle against a piece of lodestone will cause the needle to become magnetic.  Later, someone tried suspending the magnetized needle, which led to the creation of the first compass.

Geophysicists have been able to develop a mathematical model for its shape and intensity, which is called the “magnetosphere”.  It is not a perfect sphere, but instead is an imperfect sphere with many bumps and irregularities, which may be related to the complex movements of the molten iron in the outer core.   

Strong magnetic anomalies are interpreted to be caused by rocks containing magnetite, pyrrhotite, chromite, or ilmenite, while weaker anomalies may be caused by the presence of less magnetic minerals.  Mafic igneous rocks, such as gabbro, diorite or basalt, are usually the cause of the “magnetic highs”.   Other types of rocks which cause magnetic highs less frequently include skarns, and a few metamorphic rocks like greenstone (metamorphosed basalt).  Felsic igneous rocks (such as granite or rhyolite) and most sedimentary rocks are notably non-magnetic except in rare cases.  These rocks typically cause distinct magnetic lows.  Magnetic lows can be important for mineral exploration because they can be indicative of certain types of alteration.

The earth’s gravity field, like the earth’s magnetic field, is an invisible force field.  In the late 1600’s Isaac Newton demonstrated the relationship between the density (or mass) of objects and gravitational attraction between them.  He theorized the gravitational pull between two objects is inversely proportional to the square of the distance between their masses.   In mathematical terms:   F = (G)(m1)(m2) / r² ,  where F is the force of gravity, m1 and m2 are the masses of the objects, and r is the distance between the centers of the two objects, and G is the “gravitational constant”.  Any two objects have some gravitational force of attraction between them.  The amount of attraction decreases as the distance between the objects increases. 

Density is defined as mass per unit volume.  The density of a substance is directly related to the atomic weight of the element composing it, ie, elements with higher atomic weight can be thought of as being “heavier”.   Another way of stating this is that equal volumes of two different substances have different densities because each different substance has a different mass (ie, atomic weight) and crystal structure.  Measurements of mass can be made in a variety of different units (pounds, grams, etc...), likewise, measurements of volume can also be made in a variety of different units (cubic feet, cubic centimeters, etc...).  Density is most commonly measured in grams/cubic centimeter.

Minerals (or rocks) vary greatly in density, depending on their chemical make up.  Density is such a characteristic property of a substance that it may be used to identify the substance.  Geologists have found it useful to develop a system of comparing densities of different minerals or rocks.  This is done by measuring the “specific gravity”, which is essentially a comparison of the density of the substance to an equal volume of water.   For example, the specific gravity of granite is about 2.7.  This means that a cubic foot of granite, which weighs 168 pounds, is about 2.7 times heavier than a cubic foot of water, which weighs about 62.5 pounds.  Typical rock-forming, silicate minerals, such as quartz and feldspar, have  specific gravity values in the range of about 2.6 to 2.8 (Table 14 – 2).  Specific gravity values of sulfide minerals range from about 5 (pyrite) to 7.5 (galena).  Native metals (gold, platinum, etc...) have very high specific gravities ranging from about 15 to 22. 

Rock Type	Specific Gravity	Mineral	Specific Gravity
Coal	1.2 – 1.5	Sphalerite	3.8 – 4.2
Chalk	1.9 – 2.1	Chalcopyrite	4.1 – 4.3
Salt	2.1 – 2.4	Pyrrhotite	4.4 – 4.7
Serpentinite	2.5 – 2.6	Chromite	4.5 – 4.8
Granite	2.5 – 2.7	Pyrite	4.9 – 5.2
Quartzite	2.6 – 2.7	Hematite	5.0 – 5.2
Limestone	2.6 – 2.7	Magnetite	5.1 – 5.3
Gneiss	2.65 – 2.75	Galena	7.3 – 7.7
Basalt	2.7 – 3.1
Gabbro	2.7 – 3.3
Peridotite	3.1 – 3.4

The standard method of measuring the force of the earth’s gravitational field is to measure the acceleration due to gravity, which was defined by Isaac Newton:  g = (G)(m₁) / r²  (where g is the acceleration due to gravity), and F = (m₂)(g).  What the formula implies is that an object which is dropped from some height accelerates (increases its velocity) as it falls.  The acceleration can be calculated by measuring the velocity at two different times during the fall.  Likewise, the gravitational force, or gravity field, can be calculated at any specific location on the earth using the same principle.  The  value of the gravity field (acceleration) is directly related to the mass (density) of the earth beneath the station where the measurement is made.  The acceleration is measured with an instrument called a “gravimeter”.  A gravimeter measures the acceleration by sensing the pull by the earth’s gravitational field on a mass suspended from a very sensitive spring.  Gravity measurements made anywhere on the earth vary by only a few percent.  Gravity surveys use the “milligal” or “mgal” (=0.0001 gal.) as the standard unit of measure (named after Galileo).  The acceleration for one “gal” is equal to 1 cm per second per second. 

Gravimeters are used in mineral or petroleum exploration for irregularities in the predicted model of the earth’s gravity field. The gravimeter measures very tiny increases in acceleration, which suggest the presence dense rocks or minerals (such as sulfides or other dense minerals) in the subsurface (Figure 14 – 2).  The values can be plotted either along a profile or on a map (Figure 14 – 3).  Anomalous gravity highs may indicate where basement rocks are closer to the surface, or where fold structures (which may form oil traps) are located in the subsurface. 

Terrain Correction:   If a measurement is made at the base of a hill, the mass of the portion of the hill situated  topographically above the station causes an upward pull due to the attraction of the mass of the hill.  Since this counteracts the pull downward by the gravitational field, a negative correction must be made.  Likewise, if a measurement is made adjacent to a depression such as a large valley, a positive correction must be made.

Electrical methods are generally referred to as “resistivity surveys”.  Metallic minerals are relatively good conductors of electricity.  In contrast, common rock forming minerals are generally poor conductors.  This fact is the basis for geophysical exploration methods which measure conductivity to evaluate the metal content of rocks.  The methods also provide some limited information about the geometry of the subsurface metallic mineralization.  Surface electrical methods are limited to shallow depths (<500 feet), but the electrical properties of rocks can be measured at much greater depths by using special instruments sent down deep drill holes. 

“Active” electrical methods are those which introduce an artificial electrical field into the ground.  These methods utilize two electrodes placed in the ground and charged to create an electrical current which passes through the ground between them.  The resistance caused by the ground is measured and used to give an indication of the metal content.  “Passive” methods measure current flow related to ‘natural’ electrical currents.  Passive methods measure the electric potentials which develop due to the electrochemical action between minerals and pore fluids.

Native metals, metallic sulfide minerals and graphite are the best mineral conductors (Table 14 – 3).   Rocks containing abundant pore waters are also excellent conductors, in fact without these pore waters, resistivity methods would not be possible.  In general, the abundance and chemical composition of pore waters have a greater influence on conductivity than do metallic mineral grains.  For example, pore waters containing salts (sodium chloride, etc...) are the best conductors of all.  Clay minerals containing slight amounts of moisture are also excellent conductors. 

Common Rocks/Materials	Resistivity (ohm meters)	Ore Minerals	Resistivity (ohm meters)
Clay	1 – 100	Pyrrhotite	0.001 – 0.01
Graphitic Schist	10 – 500	Galena	0.001 – 100
Topsoil	50 – 100	Cassiterite	0.001 – 10,000
Gravel	100 – 600	Chalcopyrite	0.005 – 0.1
Weathered Bedrock	100 – 1000	Pyrite	0.01 – 100
Gabbro	100 – 500,000	Magnetite	0.01 – 1,000
Sandstone	200 – 8,000	Hematite	0.01 – 1,000,000
Granite	200 – 100,000	Sphalerite	1000 – 1,000,000
Basalt	200 – 100,000
Limestone	500 – 10,000
Slate	500 – 500,000
Quartzite	500 – 800,000
Greenstone	500 – 200,000

Table 14 – 3.  Resistivity of common rocks and minerals.

Figure 14 – 4.  Geometry of current flow lines and equipotential lines in a vertical section below the surface for voltage generated at stations A and B (from Dobrin, 1976).

In practice, several different pairs of electrodes are set up at different spacings, called “d-spacings”.  As the spacing between the electrode pairs increases, the detection depth increases.  In this manner, changes in resistivity with depth can be plotted on a type of cross section called a psuedosection.  This provides a means of mapping out zones of high or low resistivities.  Zones with distinctly high resistivity (low conductivity) may correspond to areas containing abundant quartz or silicification, such as vein deposits.  Zones with distinctly low resistivity may correspond to zones containing abundant metallic sulfides.  The frequency is also varied during the survey.  If metallic minerals are present, the resistivity does not change as the frequency is varied.  If metallic minerals are present, the resistivity varies dramatically as the frequency is varied.

I.P. surveys are a special type of resistivity survey.  Current flowing through the ground causes some rocks to become electrically polarized or “charged” the way a car battery is charged by an alternator when the car is running.  Chargeability is measured by sending a pulsating current through the ground.  This creates an exchange of ions between the mineral grains and surrounding pore fluids.  The exchange actually creates a voltage which acts as a barrier to current flow.  Extra voltage, called “overvoltage” is necessary to drive the current through the barrier.  When the current supply is suddenly stopped, the voltage drops immediately to an intermediate value, and then gradually dissipates.  The behavior is called “induced polarization” or “I.P.”.  Chargeability is related to the ratio of the overvoltage value to the intermediate voltage value, and the lapse time it takes the ground to “uncharge” after the electrical charge is cut off at the transmitter. 

Most modern electrical surveys measure both the resistivity and the chargeability (or “I.P. effect”).  The same equipment and electrode configuration can be used for both types of surveys.  There are at least eight different electrode configurations (or arrays) in use.  High chargeability is usually caused by the presence of metallic or conductive minerals, in contrast to high conductivity, which can be caused be salty pore fluids.  During the 1960’s, the I.P. method was used extensively to search for disseminated sulfide ores (particularly porphyry copper deposits) because of its extreme sensitivity to the presence of low grade disseminated mineralization.

Figure 14 – 5.  Psuedo-section showing chargeability values and “pants leg” shaped anomaly typically associated with a shallow conductor (from Milsom, 1996).