Gubbins D., Herrero-Bervera E. Encyclopedia of Geomagnetism and Paleomagnetism

Подождите немного. Документ загружается.

the first-order terms are considered, a symmetric matrix can be

assumed for constant H

IRM

(Cox and Doell, 1967; Stephenson et al.,

1986; Jackson, 1991). Saturation IRM also shows a grain-size depen-

dence for (titano-)magnetite, hematite, and pyrrhotite (Figure M56b).

Measurement of the anisotropy of magnetic remanence

All methods that measure the AMR require a laboratory magnetization

to be imparted along a set of known directions. The most commonly

applied laboratory magnetization is an ARM, although several studies

have used other types of magnetization, such as IRM and TRM. Since

the remanent anisotropy is a symmetric tensor of second rank, at least

six independent measurements must be made to define the tensor.

Many measurement schemes for the AMS have been proposed over

the years (see Borradaile and Henry, 1997), based on early suggestions

of Nye (1957), who suggested 13 positions (Figure M57, excluding

position 5 and 9), and Girdler (1961), who suggested nine positions

(positions 1–4, 6–8, 10, 11). Cox and Doell (1967) first described a

method to measure the anisotropy of IRM (AIRM), using 15 indepen-

dent positions (1–11, with positions 1, 3, 5, and 7 measured twice).

McCabe et al. (1985) proposed using ARM to define the remanent

anisotropy; they used the same positions as used by Girdler (1961)

to measure AMS.

Independent of the type of remanence that is applied and the number

of positions that are used, a general procedure is followed. First a sam-

ple is demagnetized and its remanent magnetization is measured as a

basis for the “demagnetized” or natural remanent state. A laboratory

magnetization is then applied along the first axis, the imparted rema-

nent magnetization is measured, and the component due to the

imparted magnetization is found by vector subtraction of the natural

remanence. The sample is again demagnetized to remove the imparted

magnetization and the demagnetized state is measured as a control.

This procedure is repeated along all the selected field positions, and

the measurements are used to define the matrix k

. Differences

between the best-fit ellipsoid and the individual measurements give

the residuals, and these can be minimized by using a least-squares

computation. Jackson (1991) proposed that a goodness-of-fit (GoF)

can be determined from

GoF ¼ RMS=ðk

 k

Þ100%

where RMS is the root mean squared residuals normalized to k

mean

The expression is an estimate of how well the anisotropy ellipsoid is

resolved. The ellipsoid can be considered well resolved for GoF values

under 10%; typically values are on the order of a few percent.

Stephenson et al . (1986) suggested an alternative method to obtain

the ellipsoid of anisotropy of remanence. They measured the individual

diagonal and off-diagonal matrix elements directly. Here a laboratory

remanent magnetization is applied along the sample x-axis and the

three components of the magnetization are measured. The sample is

demagnetized and the same field is next applied along the sample y-axis

and all the three components of the magnetization are measured.

After demagnetization the field is now applied along the sample

z-axis and the magnetization is measured. This yields nine values

that overdefine the tensor and yield a measure of the precision of

fit. This method is probably reliable for IRM, since the imparted

remanence is relatively large, allowing for accurate determination

of the magnetization components oriented normal to the applied field

direction. It is less reliable for ARM and TRM where imparted

remanences are relatively weak.

Anisotropy of anhysteretic remanent magnetization

(AARM)

The anisotropy of ARM has also been called the anisotropy of anhys-

teretic susceptibility (AAS), since normalizing the magnetization by

the field strength converts it to a pseudosusceptibility. An ARM is

acquired from the DC bias field, often on the order of 0.05–0.1 mT,

superimposed on the AC field. The strength of the bias field is on

Figure M56 (a) Variation in the intensity of anhysteretic remanent

susceptibility (k

, intensity of ARM normalized for the applied DC

field) as a function of grain size in magnetite. (b) Variation in

intensity of saturation IRM in different magnetic minerals as a

function of grain size (after Jackson, 1991).

Figure M57 Lower hemisphere, equal area plot that shows

the directions of magnetic field used to investigate the acquisition

of anisotropic remanent magnetization.

536 MAGNETIC REMANENCE, ANISOTROPY

the order of one to two times that of the Earth ’s magnetic field. The

intensity of the AC field can be chosen as desired, but is generally less

than 300 mT, due to limits of commercially available AF demagnetiz-

ing units. This limitation means that an ARM can only be imparted to

minerals with low coercivity, such as (titano-)magnetite or pyrrhotite.

A detailed description of the measurement procedure is outlined above

and has been given in McCabe et al. (1985).

Since the coercivity of ferromagnetic minerals, particularly magne-

tite and titanomagnetite, is dependent on grain size, it is possible to

impart an ARM to a particular grain-size fraction during demagnetiza-

tion in the DC bias field. This is accomplished by switching on the DC

bias field in a specific AC field range or coercivity window during

demagnetization. The imparted magnetization is known as a partial

ARM (pARM). Jackson et al. (1988) demonstrate d that the method

can be used to distinguish if a specific grain-size fraction is responsible

for the observed AARM. For example, a high coercivity window (e.g.,

80– 150 mT) will only magnetize single-domain magnetite, whereas a

low coercivity window (e.g., 0– 30 mT) would magnetize larger multi-

domain grains (see “Geological Applic ations ” sections for a practical

example).

Anisot ropy of isothermal reman ent magnet ization

(AIRM )

The advantage of AIRM is that a relatively strong DC field is applied

to the sample, such that higher coercivity phases may also be magne-

tized. As stated above, IRM is not linearly related to the applie d field;

however, a symmetric matrix can be assumed if only the first-order

terms are considered. Jelinek (1996 ) proposed a nonlinear AIRM

model that more precise ly describes the magnetization phenomena,

in which the anisotropy is also described by a symmetric second-order

tensor. The typical strength of the applied field is between 5 and

60 mT (Tarling and Hrou da, 1993). Cox and Doell (1967), who pro-

posed the method originally, used an applied field of 700 mT, which

is above the saturation of the ferromagnetic phases. The measurement

procedure is the same as outlined above. The main difficulty with their

method is the complete demagnetization between each imparted mag-

netization step. If lower fields are used, it should be possible to demag-

netize with alternating field (AF) demagnetization. Cox and Doell

(1967) rotated their samples in the 700 mT field of an electromagnet

to randomize the magnetization, since complete randomization of mag-

netic domains is reversible.

As with AARM, Cox and Doell (1967) suggested that it is possible

to examine a partial IRM. In their case, they proceeded as follows.

After measuring the sample magnetized in a 700 mT field, the sample

can be subjected to AF demagnetization in a significantly smaller field,

e.g., 100 mT. The remaining magnetization is then remeasured. The

sample can then be further demagnetized in a stronger field, e.g.,

200 mT, and then remeasured. In this manner a set of deviation vectors

is produced,

D ¼ J

 J

H = H

where J

is the measured remanent magnetization, which is produced

in the field H, and J

is the mean amplitude of a set of J

vectors aver-

aged over all the results from the different directions of the applied

field. This gives a set of D (H

) for each value of the alternating field

demagnetization (H

Several cautionary notes should be made concerning the method

and the magnetization process in AIRM. The first cautio n is with

regard to the accuracy with which a field strength can be applied.

A field strength between 5 and 60 mT is in a range where the magne-

tization of a sample is quickly acquired for (titano-)magnetite and iron

sulfides (Figure M58 ). Slight variations in field strength may lead

to apparent anisotropy that is not related to the preferred orientation

of ferromagnetic grains. Figure M58 shows a typical IRM acquisi-

tion curve for a pelagic limestone. In this example a variation in

IRM intensity on the order of 2%, comparable to the required resolu-

tion in the study of AMR, would occur if the field strength varied

by only 2mT.

Caution must also be exercised when measuring the AIRM in pyr-

rhotite-bearing rocks. A study by de Wall and Worm (1993 ) demon-

strated that the shape of the ellipsoid and degree of anisotropy are

dependent on the applied field in rocks containing pyrrhotite.

Although the orientations of the principal axes are not affected, they

showed that the degree of anisotropy in a sample of isometric pyrrho-

tite ore varied between 3.8 and 1.05 for applied field strengths between

1.5 and 100 mT, respectively. The authors attributed this effect to the

control of the observed anisotropy by magnetocrystalline anisotropy in

low applied fields, and by magnetostatic anisotropy, which arises from

the cylindrical sample shape, in strong fields. Jackson and Borradaile

(1991) reported a similar dependence of AIRM and field strength in

hematite-bearing slates, but attributed the effect to different coercivity

fractions in the rocks.

It has been noted by Tauxe et al. (1990) that following repeated

application of an IRM, the coercivity of the specimen increas ed after

exposure to the first high field, so that in spite of demagnetization to

remove the previous magnetization, some grains affected by the first

field application were no longer affected by subsequent field applica-

tions. The authors attributed this effect to hematite grains with meta-

stable domains, which change domain state during the AIRM

experiment.

Ani sotropy of thermal remanen t magnet ization (ATRM)

Heating a sample above the Curie or Néel temperature of its ferromag-

netic phases and cooling it in an applied field will produce a TRM. As

outlined in the general procedure, the sample must be first demagne-

tized. This can be achieved either with AF demagnetization or by heat-

ing the rock above its Curie temperature (T

) or Néel temperature

), and cooling in zero field. The rock is then magnetized in the first

field direction by reheating above T

or T

and cooling in a field of

fixed intensity and rate of cooling. It is important that the field intensity

and rate of cooling is the same in all heating runs. After measurement of

the imparted TRM, the sample is again demagnetized. At least six inde-

pendent heating steps are necessary to define the tensor, 12 if the sample

cannot be demagnetized in alternating fields and must be reheated above

the blocking temperature. It is important to control that no chemical

alteration has occurred, which could produce new ferromagnetic miner-

als and affect the acquired TRM. Often a thermomagnetic curve is

made on a small piece of the sample to check that the ferromagnetic

mineralogy is thermally stable.

Figure M58 Example of an IRM acquisition curve for a magnetite-

bearing, pelagic limestone.

MAGNETIC REMANENCE, ANISOTROPY 537

As for a pARM, a partial TRM (pTRM) can be acquired if the field

applied during cooling is turned on only when cooling through a par-

ticular temperature range. This requires accurate monitoring of the

temperature during cooling.

Geological applications

Several examples are presented below to illustrate how AMR can dif-

fer from AMS and the type of geologic information that can be

obtained from remanent fabrics. The most common method used in

examining AMR is AARM. The experimental method is relatively

simple and this magnetization method avoids many problems that have

been found in applying IRM or TRM. The method is suitable for mag-

netite- and possibly pyrrhotite-bearing rocks, but not for high coerciv-

ity phases, such as hematite. In this case, AIRM and ATRM are more

suitable methods. An example is given for AMR in hematite-bearing

rocks. In the examples below, the magnitudes of the principal axes

of the anisotropy ellipsoids are defined as k

 k

AMS and AARM

If the same processes leading to grain orientation affect the ferromag-

netic and paramagnetic minerals in a rock, one would expect that the

magnetic fabric of both components should be similar. The orientation

of the AARM ellipsoid is close to the AMS ellipsoid, and the main dif-

ference between the two fabrics is found in terms of the degree of ani-

sotropy, lineation, and foliation. If the two mineral components formed

at separate times in the history of the rock or if deformation mechan-

isms acting on the different mineral components varied, then it is pos-

sible that the susceptibility and remanence magnetic fabrics could be

different.

The first example illustrates how rock rheology can lead to two dif-

ferent magnetic fabrics. The Lower Paleozoic stratigraphic sequence

from the Central Appalachian Fold and Thrust Belt includes the Lower

Ordovician Coburn limestone, which is overlain by the Reedsville/

Martinsburg shale in central Pennsylvania, USA. Both lithologies

underwent the same deformation during the Alleghenian orogeny.

The AMS was measured with an Agico KLY-2 susceptibility bridge,

and the susceptibility fabric in the limestones can be characterized

by a prolate ellipsoid with k

subparallel to pencil structure in the over-

lying shales, which is also the structural trend of the major fold structures,

and k

distributed in a girdle normal to the k

axes (Figure M59).

The prolate fabric is the result of a horizontal tectonic compaction super-

imposed on the vertical bedding compaction. The shales, on the other

hand, have oblate AMS ellipsoids with k

subparallel to the pole to bed-

ding and k

aligned within the bedding plane, subparallel to the pencil

structure or fold axis. The bedding compaction controls the susceptibility

fabric in the shales.

Magnetite is the major ferromagnetic phase in both lithologies. Ani-

sotropy of ARM was investigated using a 0.1 mT DC bias field with a

150 mT alternating field, and the measurement procedure of McCabe

et al. (1985). The AARM of the limestones is similar to the AMS fab-

ric both in the orientation of the principal axes and the shape of the

ellipsoid (Figure M59). This suggests that the susceptibility and rema-

nent fabrics in the limestone are both controlled by magnetite. The

AARM in the shales, however, is characterized by prolate ellipsoids,

in which k

is well-grouped and subparallel to the pencil directions

and k

are distributed in a girdle in a plane with k

as its pole (Fig-

ure M59). This remanent fabric shows that the magnetite grains have

been more strongly affected by the horizontal compaction. The differ-

ence between the AMS and AARM fabrics can be explained by the fact

that paramagnetic clays and phyllosilicates dominate the AMS fabric,

which is largely controlled by bedding compaction. The magnetite

grains acted as rigid particles in a passive matrix during deformation,

and were therefore quicker to respond to the horizontal compression.

In the second example the AARM was determined in the Ordovi-

cian Martinsburg Formation from Lehigh Gap in eastern Pennsylvania

using ARMs applied in two different coercivity ranges. Samples were

taken from the part of the outcrop in which pencil structures are well

developed and a bedding and incipient cleavage can be identified.

Magnetite and pyrrhotite are the main ferromagnetic phases in the

rocks, as determined from IRM acquisition and thermal demagnetiza-

tion of a multiple component IRM. The AARM was first determined

using a 0.035 mT DC bias field with an alternating field of 30 mT,

and then using a 0.1 mT DC bias field with an alternating field of

100 mT, in which the DC bias field was turned on between 60 and

100 mT. The first application would only affect coarser ferromagnetic

grains, whereas the higher coercivity range would selectively magne-

tize finer grain sizes (Figure M56a). Figure M60a shows that the low

Figure M60 Lower hemisphere, equal area plots of the AARM for

(a) low-coercivity range (0.035 mT bias field, 0–30 mT AF), and

(b) high-coercivity range (0.1 mT bias field, 60–100 mT AF). The

bedding plane is shown with the dotted line and the cleavage

plane with the dashed line.

Figure M59 Lower hemisphere, equal area plot showing the

orientation of the principal axes of the AMS (upper plots) and

AARM (lower plots) for the Coburn limestone (left) and Reedsville/

Martinsburg shale (right). The principal axes k

are squares, k

are

triangles, and k

are circles in this and subsequent figures.

Bedding plane in shales is shown with the dashed line.

538 MAGNETIC REMANENCE, ANISOTROPY

coercivity AARM fabric is not well defined, but the anisotropy is

loosely related to tectonic compaction. This suggests that the coarser

grain were less efficient in orienting in the strain field. The high coer-

civity fabric is well-defined and characterized by a triaxial ellipsoid, in

which the k

axes are subparallel to the pole to cleavage, and k

axes

mirror the pencil lineation.

AIRM

Tan and Kodama (2002) investigated paleomagnetic directions in the

Mississippian Mauch Chunk Formation in eastern Pennsylvania. They

found that the inclination of the ChRM was shallower than the

expected direction and used AMR to see if the ferromagnetic grains

showed preferential flattening in the bedding plane. Hematite is the

sole ferromagnetic carrier in the red beds, and the ChRM was isolated

in thermal demagnetization above 670



C. IRM acquisition curves

indicate that 85%–90% of the saturation IRM is activated in an applied

field of 1.2 T. The AMS in the rocks has a degree of foliation on the

order of 1.01 to 1.04, which would not result in a significant inclina-

tion flattening. An IRM was applied along nine directions following

the method of McCabe et al. (1985) using a 1.2 T impulse field, and

the tensor solution was found using a method outlined in Tauxe

(2002). Samples were heated to 670



C and cooled in zero field, so that

only the ChRM-bearing hematite retains an IRM. In this manner the mag-

netic fabric due to the grains that are responsible for the ChRM can be iso-

lated. The AIRM shows that the k

axes are well-grouped along the pole to

bedding in both the undemagnetized and thermally demagnetized samples

(Figure M61). The k

and k

axes lie in the bedding plane and are slightly

better grouped in the undemagnetized samples. Both sets of data show the

same degree of lineation, but the demagnetized samples show a slightly

higher degree of foliation (Figure M61). In both AIRM cases the degree

of flattening is almost an order of magnitude higher than in the AMS.

Tan and Kodama (2002) used the AIRM and AMS data to obtain an indi-

vidual particle anisotropy factor, which was then used to correct the paleo-

magnetic inclination. They could show that, after correction, the expected

inclination was in good agreement with a European igneous paleopole of

the same age.

AMS, AARM , and ATRM

Selkin et al. (2000) made a combined study of AMS, ATRM, and

AARM in an anorthositic layer from the Upper Banded Series of the

Precambrian Stillwater complex. The susceptibility fabric, which was

measured with a KLY-2 susceptibility bridge, is compatible with the

preferred orientation of plagioclase crystals in the samples, where k

is normal to the foliation plane of the rock and k

is close to the c-axis

orientation of plagioclase (Figure M62a). Thermal demagnetization

data indicates that titanium-poor magnetite is the main ferromagnetic

phase in the samples. The AARM was measured in a 0.05 mT DC bias

field and 180 mT alternating field using six different positions. The

ATRM was determined by heating the samples to 600



C and cooling

in a 25 mT field in air; six positions were used to define the tensor.

The two different remanent fabrics are similar, but distinctly different

from the susceptibility fabric (Figure M62b,c). The AARM and ATRM

ellipsoids are prolate with k

well-grouped, coinciding with the pole to

the foliation plane defined by the plagioclase crystals and k

constrained

in the direction of the plagioclase c-axes. The authors interpret these

results as showing that single-domain magnetite grains are responsible

for the remanent anisotropy. Needles of single-domain magnetite have

been identified by TEM in plagioclase in various intrusions, including

Figure M61 Lower hemisphere, equal area plots of the AIRM for

undemagnetized samples (upper left) and thermally demagnetized

samples (upper right) from red beds of the Mauch Chunk

Formation. Principal axes ratios (lower plot) for the

undemagnetized (solid circles) and demagnetized (open circles)

samples (after Tan and Kodama, 2002).

Figure M62 Lower hemisphere, equal area plot for the magnetic fabric of an anorthosite sample, measured by (a) AMS, (b) ATRM,

and (c) AARM. Large black symbols show the principal axes of the plagioclase fabric, a nd small black symbols are the site mean

averages of the individual sample measurements (open symbols). P is the degree of anisotropy (k

), L is th e degr ee of lineation

), and F is the degree of foliation (k

)(afterSelkinet al., 2000).

MAGNETIC REMANENCE, ANISOTROPY 539

the Stillwater complex, and often the needles are perpendicular to the

plagioclase c-axes. Further support for this interpretation was obtained

from hysteresis loops made along the direction of the principal axes

of the remanence ellipsoids. Single-domain magnetite could also be

responsible for an inverted magnetic fabric in susceptibility anisotropy.

However, the same fabric would be obtained if paramagnetic minerals

are responsible for the AMS. The AMR was used to correct estimates

of paleointensity, which were affected by the preferential alignment of

the magnetite grains.

Ann M. Hirt

Bibliography

Borradaile, G.J., and Henry, B., 1997. Tectonic applications of mag-

netic susceptibility and its anisotropy. Earth-Science Reviews, 42:

49–93.

Cox, A., and Doell, R.R., 1967. Measurement of high-coercivity

magnetic anisotropy. In Collinson, D.W., Creer, K.M., and

Runcorn, S.K. (eds.), Methods in Paleomagnetism. Amsterdam:

Elsevier, pp. 477–482.

de Wall, H., and Worm, H.-U., 1993. Field dependence of magnetic

anisotropy in pyrrhotite: effects of texture and grain shape. Physics

of the Earth and Planetary Interiors, 76: 137–149.

Fuller, M.D., 1960. Anisotropy of susceptibility and the natural rema-

nent magnetization of some Welsh slates. Nature, 186: 790–792.

Girdler, R.W., 1961. The measurement and computation of anisotropy

of magnetic susceptibility of rocks. Geophysical Journal of the

Royal Astronomical Society, 5:34–45.

Graham, J.W., 1954. Magnetic susceptibility anisotropy, an un-

exploited petrofabric element. Bulletin of the Geological Society

of America, 65: 1257–1258.

Hargraves, R.B., 1959. Magnetic anisotropy and remanent magnetiza-

tion in hemo-ilmenite ore deposits at Allard Lake, Quebec. Journal

of Geophysical Research, 64: 1565–1578.

Howell, L.G., Martinez, J.D., and Statham, E.H., 1958. Some observa-

tions on rock magnetism. Geophysics, 23: 285–298.

Jackson, M., 1991. Anisotropy of magnetic remanence: a brief review

of mineralogical sources, physical origins, and geological applica-

tions, and comparison with susceptibility anisotropy. Pure and

Applied Geophysics, 136:1–28.

Jackson, M.J., and Borradaile, G., 1991. On the origin of the magnetic

fabric in purple Cambrian slates of north Wales. Tectonophysics,

194:49–58.

Jackson, M., Gruber, W., Marvin, J., and Banerjee, S.K., 1988. Partial

anhysteretic remanence and its anisotropy: Applications and grain

size-dependence. Geophysical Research Letters, 15: 440–443.

Jelinek, V., 1996. Theory and measurement of the anisotropy of iso-

thermal remanent magnetization of rocks. Travaux Géophysique,

37: 124–

134.

McCabe, C., Jackson, M., and Ellwood, B.B., 1985. Magnetic aniso-

tropy in the Trenton limestone: results of a new technique,

anisotropy of anhysteretic susceptibility. Geophysical Research

Letters, 12: 333–336.

Nye, J.F., 1957. Physical Properties of Crystals. London: Oxford

University Press.

Rochette, P., Jackson, M.J., and Aubourg, C., 1992. Rock magnetism

and the interpretation of anisotropy of magnetic susceptibility.

Reviews of Geophysics, 30: 209–226.

Selkin, P.A., Gee, J.S., Tauxe, L., Meurer, W.P., and Newell, A.J.,

2000. The effect of remanence anisotropy on paleointensity esti-

mates: a case study from the Archean Stillwater Complex. Earth

and Planetary Science Letters, 183: 403–416.

Stacey, F.D., and Banerjee, S.K., 1974. The Physical Principles of Rock

Magnetism. Amsterdam: Elsevier Scientific Publishing Company.

Stephenson, A., Sadikun, S., and Potter, D.K., 1986. A theoretical

and experimental comparison of the anisotropies of magnetic

susceptibility and remanence in rocks and minerals. Geophysical

Journal of the Royal Astronomical Society, 84: 185–200.

Tauxe, L., Constable, C., Stokking, L., and Badgley, C., 1990. Use of

anisotropy to determine the origin of characteristic remanent mag-

netization in the Siwalik red beds of northern Pakistan. Journal of

Geophysical Research—Solid Earth and Planets, 95: 4391–4404.

Tan, X.D., and Kodama, K.P., 2002. Magnetic anisotropy and paleo-

magnetic inclination shallowing in red beds: evidence from the

Mississippian Mauch Chunk Formation, Pennsylvania. Journal of

Geophysical Research—Solid Earth, 107B11):2311, doi: 10.1029/

2001JB001636.

Tarling, D.H., and Hrouda, F., 1993. The Magnetic Anisotropy of

Rocks. London: Chapman & Hall.

Tauxe, L., 2002. Paleomagnetic Principles and Practice. Dordrecht:

Kluwer Academic Press.

Cross-references

Magnetic Susceptibility, Anisotropy

Magnetization, Anhysteretic Remanent (ARM)

Magnetization, Isothermal Remanent (IRM)

Magnetization, Natural Remanent (NRM)

Magnetization, Thermoremanent (TRM)

MAGNETIC SHIELDING

There are numerous uses for magnetic shields in research and consumer

electronics. In geophysical research, magnetic shields are used in

paleomagnetic laboratories to protect samples and improve the perfor-

mance of sample magnetometers. Research shields range in size from

10 cm to 15 m.

The irony in magnetic shielding theory is that a shield actually adds

another magnetic field. This new field is specifically designed to be

equal and opposite to the existing field, resulting in a nullification

of, not a removal of the ambient magnetic fields. This is analogous

to gravitational forces at Lagrangian points, or the balance of centripe-

tal and gravitational forces on orbiting bodies, or noise cancellation by

antiphase headphones.

Materials used

Diverse materials and techniques are used for magnetic shielding. The

choice depends on the kind of fields to be opposed (e.g., AC vs DC)

and the specific application. In general, shielding can be accomplished

by using one of two physical phenomenon, either electrical conductiv-

ity or strong magnetic responses. Electrical conductors carry currents

that produce a magnetic field perpendicular to the current direction.

To generate these currents in a conductive sheet, there must be an ori-

ginal, varying magnetic field. Within the conductor, these currents pro-

duce a new magnetic field in the opposite direction to this exciting

field ( Lenz’s law). By surrounding a site with a good conductor, the

volume will be shielded from varying fields coming from all direc-

tions. The important parameters for efficient shielding by conductors

are: conductivity and thickness of the shielding material, frequency

of the magnetic field, and the completeness of enclosure. Of course,

electrical wires can be powered externally to produce a magnetic field

of almost any configuration, including the nullification case of shield-

ing. This will be considered in detail later.

Materials with strong magnetic properties, usually metals containing

Fe, Ni, or Co, can react to changing magnetic fields by generating an

oppositely directed field (magnetic permeability) or can be tuned to

oppose a steady field (magnetic remanence). Efficient shielding by mag-

netic materials depends on the material, its treatment history, as well

as the ambient field characteristics of frequency, strength, gradient,

and source direction.

540 MAGNETIC SHIELDING

Sources of magnetic fields

The goal of magnetic shielding can be either to nullify ambient mag-

netic fields (external source shielding, described in detail later) or to

nullify local, instrument-generated magnetic fields (internal source

shielding). Many consumer electronics and medical diagnostic instru-

ments generate large local magnetic fields. Here the goal of shielding

is to protect other circuits and conductors within the instruments from

unwanted induced currents, and to allow the normal functioning of

other, nearby electronic equipment. The larger the source field, the

more important shielding becomes. Significant care and expense is

taken to shield the immediate environment from monitors and TVs

(cathode-ray tube types), and nuclear magnetic resonance (NMR)

instruments. Since these internal sources generate rapidly varying

magnetic fields, the shielding material must have either high magnetic

permeability or high conductivity.

External source shielding has the challenge of encountering a differ-

ent magnetic environment (differences in frequency, strength, and

direction) in each location. In theory, the entire range of the electro-

magnetic spectrum could be shielded against, but an equally broad

range of shielding materials would need to be used simultaneously.

Great design simplifications can be realized by restricting the shielding

to a specific frequency or a narrow range of frequencies.

Shielding in paleomagnetic research

Paleomagnetic research uses shields to protect samples during rema-

nence measurements and during demagnetization experiments. In

general, shielding design focuses on the reduction of DC fields (stable

in the time domain). Most laboratory sites have large field gradients

owing to the steel in pipes, ducts, building framework, furniture, and

research equipment. Where space permits, some laboratories have fab-

ricated specific, isolated buildings without steel components, to create

a low-gradient environment. Isolation from varying field sources is

another critical design feature. These sources can be: vehicles (moving

or parked), elevators (especially piston types), steel doors, steel con-

tainers, mass spectrometers, magnetic mineral separators, circuit panel

boxes, and electric-powered mass transit.

Three different shielding techniques are used in paleomagnetic

laboratories: coils, high-permeability magnetic sheets, and remanent

magnetic sheets. Current flow in coils of wire can be adjusted to pro-

duce any magnetic field desired, including a field that matches, but is

in the opposite direction to, a component of the ambient field. Sets of

coils can be configured to control a component of the field over a spe-

cified volume, e.g., a pair of coils, spaced apart a distance equal to

their radius (the Helmholtz arrangement). For a review of other multi-

coil arrangements see Kirschvink (1992). The advantages of coil sys-

tems are: relatively low cost, open arrangement, flexible positioning,

and adjustable field strength (net fields other than zero). Disadvantages

are: small-shielded volume relative to apparatus size (shielded volume

to coil volume about 1:100), the need for a large DC power supply,

continuous cost for electrical power, and difficult access to the

shielded center of a three-component set.

High-permeability magnetic materials are extensively used for

research shielding. Nickel-rich metals are common, along with newly

developed materials such as glass and amorphous metals. Rolled into

cylinders, these shields can be arranged concentrically to provide excel-

lent shielding. Design considerations for these cylindrical shields are:

orientation (optimum direction of cylinder axis is perpendicular to the

field), open or closed ends (open-ended cylinders have higher fields at

least 2 diameters inside), and treatment history (both metallurgical and

magnetic degaussing). Small room-sized shields have been made from

high-permeability sheet material (Patton and Fitch, 1962). There are

hybrid designs for these larger shields that utilize low-cost remanent

materials for the bulk of the shielding. The high-permeability magnetic

materials are used only as an inner shield, taking advantage of their

ability to react to (shield against) weak and variable magnetic fields.

The advantages of high-permeability magnetic materials are: effective

shielding over a wide range of frequency and field strengths, including

DC fields. Disadvantages are: cost, annealing and processing restrictions

on size and shape, and permeability reduction if deformed.

Remanent magnetic materials have been extensively used (since

1980) for large volume, room-sized shields (Scott and Frohlich,

1985). Originally developed for the transformer industry, silicon steel

alloys are used as magnetostatic shields in many paleomagnetic labora-

tories. Magnetostatic or DC shielding incorporates a series of design

simplifications that allow the fabrication of large (15 m) and user-

friendly enclosures (open doorways and windows). Therefore, an

entire paleomagnetic laboratory along with thousands of specimens

can be contained inside a shielded enclosure. Remanent shielding

materials must be magnetically tuned (using anhysteretic remanence)

to generate the balanced opposing (or shielding) field. The advantages

of remanent magnetic materials are: low cost, wide range of size, and

high magnetic field strength. Disadvantages are: postfabrication tuning

(remagnetization) is required, and ineffective shielding against variable

magnetic fields.

One distinguishing feature of shields that use magnetic materials is

the concentric layers of shielding material, referred to as shielding

stages. While the typical two-stage shield provides a sufficient field

reduction for many applications, shields with up to six stages have

been fabricated for specific research applications. Each shielding stage

can only respond to the magnetic field where it is located, so that inter-

ior stage shields respond to much smaller fields than the initial (outer-

most) shield. Therefore, interior shields are less effective in both

absolute strength and percent reduction in residual fields. This is

caused by the reduced magnetic alignment in weaker magnetic fields,

called the Langevin alignment function. This applies to both high-per-

meability materials (permeability is much smaller in lower fields) and

in remanence shields (the smaller biasing field lowers the strength of

anhysteretic remanence). Using a three-stage shield as an example,

we start with the outer stage in the Earth’s field (50000 nT), which

can reduce the interior field by 95% (down to 2500 nT). The middle

stage is in the 2500 nT field, and can reduce the interior field by

90% (down to 250 nT). The inner stage is in the 250 nT field, and

can reduce the interior field by 80% (down to 50 nT). Most shields

are assembled starting with the outer shield, progressively adding

stages to the interior. The residual interior field provided by a magnetic

shield is proportional to both the initial strength of the ambient mag-

netic field, and to the magnitude of pre-existing field gradients.

The spacing between shielding stages is approximately 10% of the

shield diameter. Why so far apart? And why not fill the interval

between the stages with shielding material? Wills (1899) examined

these questions in detail for both the two- and three-stage cases. Coun-

terintuitively, his numerical (2D) models showed that for two-stage

shields the residual field was five times higher if the area between

the shielding stages was solid. Theoretically, this decreased shielding

with increased mass is analogous to the decreased field from a sin-

gle-domain particle as it increases in volume to a multidomain state.

Thick magnetic material will magnetically interact with itself, in a

process of self-cancellation that produces a small external field. Like-

wise, the spacing between stages allows each stage to be magnetically

independent, instead of being linked by dipolar interaction into a lower

energy state. Wide spacing between stages is also an important design

feature when large field gradients are present.

If the goal is to shield against high-frequency fields ( >100 Hz), then

thick electrically conductive sheets are used. Thickness is an advan-

tage in this case, where the conductivity phenomenon of skin-depth

is important. Induced electric currents should be able to flow uninter-

rupted in any direction, thus generating the opposing magnetic field.

With conductive materials, multiple shielding stages are not as useful

as concentrating the material in a single, thick layer. At low frequen-

cies (<100 Hz), the rate of change of the magnetic field will not be

able to generate sufficient current flow to create the opposing magnetic

field. In the case of biological research, the shielding must respond to

a wide range of frequencies, therefore dual material shields are usually

MAGNETIC SHIELDING 541

fabricated. One of the design challenges is that these two materials react

orthogonally. Conductive material react to magnetic fields that arrive per-

pendicular to the sheet, as compared to magnetic material that react to

magnetic fields that arrive parallel to the sheet.

Gary R. Scott

Bibliography

Kirschvink, J.L., 1992. Uniform magnetic fields and double-wrapped

coil systems: improved techniques for the design of bioelectromag-

netic experiments. Bioelectromagnetics, 13: 401–411.

Patton, B.J., and Fitch, J.L., 1962. Design of a room-size magnetic

shield. Journal of Geophysical Research, 67: 1117–1121.

Scott, G.R., and Frohlich, C., 1985. Large-volume magnetically

shielded room, a new design and material. In Kirschvink, J.L.,

Jones, D.S., and Macfadden, B.J. (eds.), Magnetite Biomineraliza-

tion and Magnetoreception in Organisms. New York: Plenum

Press, pp. 197–220.

Wills, A.P., 1899. On the magnetic shielding effect of trilamellar sphe-

rical and cylindrical shells. Physical Review, 9: 193–213.

MAGNETIC SURVEYS, MARINE

Definition

A marine magnetic survey is the measurement of Earth’s magnetic

field intensity or its components (such as vertical component) along

a series of profiles over an area of interest with the objective of mea-

suring the magnetism of the ocean floor. The seafloor is known to

be highly magnetized in a pattern of magnetic stripes that record the

history of Earth’s magnetic field behavior, which in turn, allows mar-

ine scientists to determine the age and history of seafloor spreading in

the ocean basins (see Vine-Matthews-Morley hypothesis). Magnetic

surveys are also used to locate a concentration or absence of magnetic

minerals in the search for economic ore deposits (Sheriff, 1981). Var-

iations in the observed magnetic field from the predicted regional field

are called magnetic anomalies and are attributed to variations in the

distribution of materials having different magnetic polarity, remanent

magnetization, and/or susceptibility. The SI units of measurement for

Earth’s magnetic field intensity are nanoteslas (nT) or in the older

cgs system gammas, where 1 gamma is equal to 1 nT. Earth’s mag-

netic field intensity varies between 30000 and 60000 nT and magnetic

anomalies measured at sea are typically in the range of 1–1000 nT.

History of marine magnetic surveying

Magnetic field measurements have been made at sea since the early

days of marine exploration, although these consisted primarily of

directional measurements for declination and inclination (see History

of geomagnetism, Voyages making geomagnetic measurements,

Halley, Carnegie Institute of Terrestrial Magnetism). One of the first

observations that magnetic fields at sea may be caused by seafloor

geologic sources was made by Cole (1908) who compiled a declina-

tion chart of the western Atlantic Ocean around Bermuda in which

“sources of considerable disturbance” could be discerned. The research

vessel (R/V) Carnegie (see Carnegie research vessel) carried out the

first global marine magnetic surveys for the purpose of mapping

Earth’s magnetic field. During World War II, the rise of submarine

warfare precipitated the development of magnetic airborne detectors

(MAD); a self-orienting fluxgate magnetometer flown by an aircraft

above the ocean to detect submerged submarines (Rumbaugh and

Alldredge, 1949). This development allowed for continuous magnetic

field intensity measurements to be made at sea. After the war, one of

these magnetometers was modified into a towed marine magnetometer

(Heezen et al., 1953) and in 1948 the first shipboard magnetic anom-

aly profile was collected across the Atlantic Ocean from Dakar to

Barbados with R/V Atlantis. The profile crossed the then unknown

Mid-Atlantic Ridge and recorded a strong anomaly over the rift valley

that is now known to be the Brunhes central anomaly that marks the

locus of the active mid-ocean ridge spreading center where new crust

is being created at the seafloor. Strong magnetic lows were also found

over the flanks of the ridge, which are now known to be the reversed

crust of the Matuyama Chron and again part of the seafloor spreading

magnetic stripes (see Vine-Matthews-Morley hypothesis).

In 1954, Arthur Raff at Scripps Institution of Oceanography began

using a fluxgate magnetometer based on the Lamont design. Raff

worked with Russell Varian and Martin Packard to calibrate the

ship-towed magnetometer against their newly developed proton pre-

cession magnetometer that provided an absolute value of Earth’s mag-

netic field intensity (Glen, 1982). Proton precession magnetometers do

not suffer from drift problems or require calibration like the fluxgate

sensors and is the standard tool of today ’s magnetic surveying prac-

tice. These early surveys proved successful and Arthur Raff teamed

up with geophysicist Ronald Mason visiting from Imperial College

in London to tow their magnetometer as part of a large bathymetric

survey off the west coast of North America using the US Coast and

Geodetic Survey (USCGS) ship Pioneer. This survey was published

by Raff and Mason (1961) and became one of the key data sets leading

to the acceptance of the ideas about how the marine magnetic stripe

pattern was recorded by the ocean crust and subsequently preserved

by the seafloor spreading process (Dietz, 1961; Hess, 1962).

Today, modern marine magnetometers have a precision of between

1 and 0.1 nT using proton precession sensors, optically pumped

cesium magnetometers or more recently, the Overhauser nuclear pre-

cession magnetometer (see Observatory instrumentation). While mag-

netometers are still towed behind ships for conventional surveys,

magnetic measurements are also made from a variety of seagoing plat-

forms including deep-diving submersibles (Macdonald et al., 1983),

deep-towed sleds (Macdonald, 1977), remotely operated vehicles (ROVs)

(Johnson et al., 2002), and autonomous underwater vehicles (AUVs)

(Tivey et al., 1997). Airborne magnetic surveys are also an impor-

tant method of surveying particularly over the sediment-dominated con-

tinental shelf areas in the hunt for natural resources of economic value

such as oil and gas. The US Naval Oceanographi c Office also ran an

airborne magnetic survey program called Project MAGNET for many

years with the aim of mapping Earth’s magnetic field intensity on a

global scale (see Project MAGNET).

Methodology, analysis, and processing

Sea surface marine magnetic surveys are typically obtained by towing

a magnetometer sensor behind a ship. The sensor is towed at the end

of a cable with electrical conductors that enable the sensor to be pow-

ered and communicated to from the ship (Figure M63). The sensor is

towed between 3 and 6 ship lengths behind the ship in order to reduce

the magnetic effects of the ship on the measurement. The optimal lat-

eral line spacing for a survey is directly related to the water depth

because this is the filtering distance over which the magnetic field

has propagated. The typical depth of the ocean ranges between 2 and

6 km so that a line spacing of 2 to 10 km is usually sufficient for most

surveys (Figure M64). Magnetic surveying is often done in conjunc-

tion with other geophysical techniques such as swath bathymetric sur-

veying, gravity and seismic profiling, any of which may determine the

ultimate line spacing required. The best approach to magnetic surveying

is to collect profiles perpendicular to the magnetic “grain” or strip-

ing that characterizes the ocean basins.

The first step in processing is the removal of noise spikes that may

have been produced from external sources, such as motors and genera-

tors. At equatorial latitudes it is also usually necessary to remove the

diurnal variation in Earth’s magnetic field due to the combined effects

542 MAGNETIC SURVEYS, MARINE

of the equatorial electrojet and the weak equatorial geomagnetic field.

Diurnal records can be obtained from nearby geomagnetic observa-

tories on land or with a temporarily installed base station on the sea-

floor. The next step is to merge the magnetic time series data with

the ship’s position. In the modern era this is easily accomplished with

Global Positioning Satellite (GPS) data, but in the past this was per-

haps the largest source of error, especially in remote areas of the

ocean. Once the position of the magnetic field measurement is known

the regional magnetic field is removed to reveal the residual “magnetic

anomaly” field. Typically, the predicted regional field is obtained from

the global International Geomagnetic Reference Field (see IGRF,

DGRF) spherical harmonic model for the appropriate date and geogra-

phical location of the survey. The suitably processed data can then be

plotted as wiggle plots (Figure M65) or depending on the data density

gridded into a contour map (Figure M66).

Magnetic anomaly highs represent places where the local magnetic

field exceeds the regional field and magnetic lows are where the local

magnetic field is less than the regional field. For marine magnetic sur-

veys these high- and low-magnetic anomalies correspond to magne-

tized zones of normal and reverse polarity ocean crust, respectively.

The source of marine magnetic anomalies is primarily the volcanic

rocks that comprise ocean crust and typically not the overlying

Figure M64 Map showing the seafloor depth for a portion of the western flank of the MidAtlantic Ridge near 26



N. Bold black lines

indicate the location of the midocean ridge spreading axis and arrows indicate direction of seafloor spreading. Data from Tivey and

Tucholke (1998). Thin lines indicate shiptracks.

Figure M63/Plate 9a (a) Photograph showing a marine proton precession magnetometer and tow cable on a winch before launch.

(Photo taken by Maurice Tivey, Woods Hole Oceanographic Institution.) (b) Photograph showing the deployment of a marine proton

precession magnetometer from the stern of a research vessel. (Photo taken by Maurice Tivey, Woods Hole Oceanographic Institution.)

MAGNETIC SURVEYS, MARINE 543

sediments. The magnetic anomalies can be correlated and identified

with the predicted anomalies from the geomagnetic polarity timescale

(GPTS) and thereby allow for an isochron map of the ocean crust to be

generated (Figure M67). This allows for a history of seafloor spreading

to be documented (see Figure M68) and for spreading rates to be cal-

culated.

Magnetic anomalies vary in their morphology according to their

latitude of formation. Magnetic anomalies arise not only from the mag-

nitude of the magnetization but also the direction of the magnetization.

Unlike gravity, which is directed toward Earth’s center, magnetic field

lines originate on south poles and end on north poles. Thus, magnetic

anomalies are centered over the magnetized body at the poles, but off-

set and phase-shifted at midlatitudes with magnetic lows and highs

marking the north and south ends of a normal polarity magnetized

body. At the equator a normal polarized body produces a magnetic

low. These phase shifts can be compensated for by either reducing

the anomaly to the pole (see Reduction to pole) or by computing the

crustal magnetization (Figure M67).

Magnetic surveys are also collected with deep-towed sensors or

other moving platforms such as submersibles or AUVs (see for exam-

ple Tivey et al., 1997, 2003). Such surveys are usually detailed studies

encompassing small areas of the seafloor (<10 km). Additional

Figure M66/Plate 9b Contour map of the same data shown in Figure M65 showing the magnetic anomaly stripes that characterize the

seafloor over most of the ocean basins. Contour interval is 50 nT. Thin black lines indicate the ship’s tracklines. Map based on data

from Tivey and Tucholke (1998).

Figure M65 Map showing the magnetic anomaly wiggles along the tracklines of a marine magnetic survey on the western flank of

the MidAtlantic Ridge near 26



N. Positive anomalies are shaded gray and represent a firstorder normal polarity crust. Negative

anomalies are unshaded and represent the reverse polarity crust. Data from Tivey and Tucholke (1998).

544 MAGNETIC SURVEYS, MARINE

processing steps are required in the analysis of these kinds of data. The

depth and height above the seafloor of the vehicle or platform must

be known with some precision. Various algorithms exist to continue

the data from an uneven observation surface onto a level observation

plane above the topography (see Tivey et al., 2003). If the type of mag-

netic sensor used in the survey is a vector magnetometer rather than

an absolute sensor then the orientation of the vehicle (i.e., its attitude)

must also be known. Finally, navigation requires specialized acoustic

Figure M68/Plate 9d Map showing the interpolated magnetic polarity stripes with normal polarity indicated by black and reverse

polarity in white. Thin black lines indicate the boundaries of individual spreading ridge segments through time. The geomagnetic

polarity timescale is shown at the bottom to correlate with age.

Figure M67/Plate 9c Map of the calculated crustal magnetization based on the seafloor depth (Figure M64) and magnetic anomaly data

(Figure M66). Black dots and adjoining lines indicate the reverse polarity isochrons, whereas the white triangle symbols and adjoining

lines indicate normal polarity chrons. Selected major chrons are identified and can be converted to age using the geomagnetic polarity

timescale shown at the bottom of the figure. Map based on data from Tivey and Tucholke (1998).

MAGNETIC SURVEYS, MARINE 545