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

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Coordinated biostratigraphic and magnetostratigraphic investigations

have correlated most important paleontological stage and substage

boundaries from the Late Jurassic to the present with the corresponding

geomagnetic polarity sequences.

The reversal history of the Pleistocene and Upper Pliocene was

established in the 1960s jointly from radiometrically dated lavas and

modern deep-sea sediments (Opdyke, 1972). During the 1970s and

1980s, magnetostratigraphic research confirmed the marine magnetic

record in Paleogene and older rocks. The verification of Miocene

polarity chrons was delayed by high reversal rate, weak magnetization,

and unstable remanent magnetizations in many sections. However, coor-

dinated magnetostratigraphy, biostratigraphy, and cyclostratigraphy in

marly sections in Sicily and Crete confirmed and dated the Early Plio-

cene and Upper Miocene polarity sequence (Krijgsman et al., 1995).

Recent results from ODP sites have given excellent “fingerprint” corre-

lation of Early Miocene polarity chrons (Parés and Lanci, 2004).

Magnetostratigraphic studies in limestone sections in the Umbrian

Apennines and Southern Alps confirmed and dated geomagnetic polarity

history from the Jurassic-Cretaceous boundary to the Oligocene-Miocene

boundary (Channell and Erba, 1992; Channell et al., 1979, 2000; Lowrie

et al., 1980, 1982; Channell and Medizza, 1981; Lowrie and Alvarez,

1981, 1984). The Cretaceous-Tertiary boundary was correlated within

polarity chron C29r and the Santonian-Campanian stage boundary was

located just older than polarity chron C33r (Alvarez et al., 1977; Lowrie

and Alvarez, 1977) (Figure M164).

Occasional reports of reversed polarity zones in magnetostrati-

graphic sections have given rise to speculation about the continuity

of the 37-Ma long CNPS and the origin of the Cretaceous Quiet Zone.

However, in each instance the magnetostratigraphic results have been

associated with a rock magnetic complication, such as remagnetization

in redbeds, lack of independent verification, or possible sample inver-

sion. The paleomagnetic evidence supports the interpretation of the

CNPS as an uninterrupted interval of constant normal polarity. It

may represent a distinct behavioral regime of the geomagnetic field,

different in its nature from the preceding and following episodes of

reversing polarity.

The youngest reverse polarity chron CM0r of the M-sequence is

located just above the Barremian-Aptian boundary (Channell and

Erba, 1992). Correlation of the Jurassic-Cretaceous boundary with

the M-sequence anomalies was found to be slightly dependent on the

paleontological dating scheme used to define the boundary (Lowrie

and Channell, 1984; Ogg, 1984; Lowrie and Ogg, 1986). Magneto-

zones in the Kimmeridgian and Tithonian stages of the Upper Jurassic

were correlated in magnetostratigraphic sections in southern Spain with

M-sequence anomalies M18-M25 (Ogg et al., 1984) (Figure M165).

Many details of these correlations and polarity sequences have been

further confirmed in ODP drillcores.

In addition to the well-documented magnetic anomalies of the M-

sequence and C-sequence, the marine magnetic record contains many

short wavelength, low-amplitude magnetic anomalies (tiny wiggles).

The origin of these features is uncertain, and they have been desig-

nated as “cryptochrons.” They can be modeled equally well as very

short polarity chrons with typical durations of 10 ka or less, or

as intensity fluctuations of the paleomagnetic field (Cande and

LaBrecque, 1974; Cande and Kent, 1992b). They have generally not

been identified in detailed magnetostratigraphic investigations. Although

some may indeed represent polarity reversals, most are probably due to

paleointensity variations and are not part of the reversal sequence.

Magnetostratigraphy older than the M-sequence

Magnetostratigraphy in older rocks suffers from the absence of a mar-

ine magnetic anomaly record. Geomagnetic polarity history must be

pieced together from magnetic stratigraphy alone. This requires over-

lapping sections to confirm the polarity sequence. Until now, the

pre-M29 magnetic anomalies have not been securely and indepen-

dently established in the oceanic record. Their appearance resembles

the controversial tiny wiggles found in the Cenozoic record (Cande

and Kent, 1992b) and they may be due to paleointensity fluctuations

rather than polarity reversals. Deep-tow studies are in progress to ver-

ify the anomaly correlations. The assumed polarity sequence must then

be confirmed and dated independently by magnetostratigraphy. This is

a difficult task. Middle Jurassic limestone sections have given good

magnetostratigraphic data but the magnetozones correlate only tenu-

ously with the marine polarity record corresponding to anomalies

M29-M41 (Steiner et al., 1985, 1987).

Lower Jurassic results from the Breggia Gorge in Switzerland (Hor-

ner and Heller, 1983) showed a very well-defined magnetostratigraphy

for the Carixian (early Pliensbachian) to Bajocian stages of the Early

to Middle Jurassic (Figure M163). This record stands as the reference

magnetic polarity stratigraphy for this part of the timescale. Because of

the rareness of comparable sites with suitable magnetic properties for

magnetic polarity analysis, it has only been verified in part. Moreau

(2002) recently established the magnetostratigraphy of a drillcore from

the Paris Basin and correlated it successfully to the Pliensbachian part

of the Breggia Gorge record.

It is a great challenge for current and future magnetostratigraphic

research to piece together the history of geomagnetic polarity rever-

sals for earlier periods in the Mesozoic and Paleozoic. The magnetic

polarity stratigraphy of the Late Triassic is a focus of current activ-

ity. Excellent results have been obtained in overlapping drillcores

from continental redbeds from the Newark Basin in New Jersey,

USA (Kent et al., 1995). This study required subtle analysis of

Figure M164 Correlation of Cenozoic and Upper Cretaceous

limestone sections in the Umbrian Apennines, Italy, with the

oceanic polarity record on the basis of magnetic polarity

stratigraphy (after Lowrie and Alvarez, 1981).

666 MAGNETOSTRATIGRAPHY

thermal demagnetization data; the polarity stratigraphy was dated

by a combination of paleontology and Milankovich-related cyclos-

tratigraphy. As the Breggia Gorge is for the Lower Jurass ic, the

magnetostratigraphy of the Newark Basin is the standard for com-

parison for the Late Triassic. Confirmation by correlation with

other sections, especially of marine origin, is in progress.

Rock magnetic stratigraphy

Susceptibility stratigraphy

Although polarity stratigraphy is the best-known magnetic method

for establishing the stratigraphic v ariation in a section and for corre-

lating between stratigraphic sections, other magnetic me thods are

also possible. For example, magnetic susceptibility, remanent mag-

netizations, and coercivity parameters have all been invoked in

magnetic studies of young sediments to show small variations in

mineralogy or grain size that might have an environmental or

paleoclimatic cause.

Magnetic susceptibility is a measure of the ease with which a mag-

netization can be induced in a sample. It is determined by all the

mineral grains in a sample; that is, by the type, concentrations, and

grain sizes of any ferromagnetic minerals, as well as by the paramagnetic

and diamagnetic (i.e., nonferromagnetic) minerals in the sample. If

magnetite is present, its high intrinsic susceptibility may dominate the

susceptibility of a sample, and fluctuations in susceptibility within a

sedimentary sequence may be related to variations in magnetite content.

An example is shown in Figure M166. The loess deposits of Central

China form great thicknesses of sediment in which loess and paleosol

layers alternate with each other. The loess are windblown sediments

that form in cold, dry conditions, whereas the paleosols are thought

to form under warmer, moist conditions. The natural remanent magne-

tization of the deposits gives a well-defined magnetic polarity strati-

graphy that correlates the oldest layers with the end of the Gauss

Epoch (top of Chron C2An.1n) about 2.6 Ma ago (Heller and Liu,

1984; Heller and Evans, 1995). The magnetostratigraphy of samples

from a borehole at Luochuan in China showed that the loess deposits

were much older than previously thought. The magnetic susceptibility

correlated with the lithology, showing higher values in the paleosols

than in the loess layers. The magnetic mineral in both lithologies is

dominantly magnetite. The magnetite enhancement in the paleosols

may be related to in situ formation of a new phase of magnetite as a

result of the change to a warmer, moister climate. This interpretation

has been carried a stage further in an attempt to estimate the record

of paleoprecipitation. The existence and interpretation of magnetite

enhancement in the paleosols is still a controversial subject. Neverthe-

less, the example demonstrates the potential of magnetic measure-

ments in an environmental context.

Coercivity stratigraphy

In contrast to magnetic susceptibility, the remanent magnetization of a

sample, whether naturally acquired or artificially produced by an applied

magneticfieldinthelaboratory,iscarried only by ferromagnetic minerals.

Figure M165 Correlation of Lower Cretaceous and Upper Jurassic limestone sections with the polarity record derived from the

M-sequence of oceanic magnetic anomalies (after Lowrie, 1989).

MAGNETOSTRATIGRAPHY 667

The most common are magnetite, hematite, maghemite, and pyrrhotite.

The ability to retain a remanent magnetization is a coercivity para-

meter, and is dependent on grain size, temperature, and the applied

magnetic field. A commonly used remanent magnetization is produced

in the laboratory by applying a magnetic field to a sample and then

removing it. The residual magnetization is called the isothermal rema-

nent magnetization (IRM). Different grain-size fractions are activated

by different applied fields, and at different temperatures. For example,

Figure M166 Magnetic polarity stratigraphy and susce ptibility stratigraphy of loess and paleosol deposits in a borehole at Luochuan,

China (after Heller and Evans, 1995; courtesy of F. Heller).

Figure M167 Magnetic stratigraphy in a core from a high Alpine lake (Bachalpsee, Switzerland). The ratios of isothermal remanent

magnetizations (IRM) acquired at 290 and 77 K, and in fields of 0.15 and 1 T at room temperature, are determined by variations

in coercivity of ferromagnetic minerals. They are compared to climatic proxies that reflect organic content and vegetation changes

(after Lanci et al., 1999).

668 MAGNETOSTRATIGRAPHY

ultrafine-grained magnetite that is superparamagnetic (unable to hold a

remanent magnetization) at room temperature becomes stable at low

temperature. The IRM of a sample containing superparamagnetic

grains of magnetite is thus larger at low temperature than at room tem-

perature. The ratio of IRM acquired at room temperature (290 K) and

after immersion in liquid nitrogen at a temperature of 77 K (written

IRM

290 K

/IRM

77 K

) is an indication of the relative amounts of superpar-

amagnetic magnetite in a sediment. Similarly, the ratio of IRM acquired

in fields of 0.15 and 1.0 T (written IRM

0.15 T

/IRM

1.0 T

) is an indication

of the relative magnitude of low-coercivity magnetite and higher coer-

civity grains in a sample. For example, in a magnetostratigraphic section

of Upper Cretaceous Scaglia Rossa limestone at Gubbio, Italy, the var-

iation of this parameter showed the relative amounts of magnetite and

hematite in the section (Lowrie and Alvarez, 1977). In a sediment core

from Bachalpsee (Figure M167), a high-alpine lake in the Swiss Alps, this

ratio has a value around 0.8 at the bottom of the core. About 80% of

the IRM is carried by low-coercivity magnetite at this depth, which corre-

sponds to the Younger Dryas biozone. The remainder may be due to high-

coercivity magnetite or even harder magnetic minerals. Higher in the core,

where the ratio is almost 1.0, low-coercivity magnetite is virtually the only

ferromagnetic mineral present. The ratio decreases again upward in the

core. The parallel behavior of the IRM

290 K

/IRM

77 K

ratio shows that

grain-size variation of the magnetite is probably the cause of both curves.

The loss of ignition at 550



C, an indicator of organic carbon

content, parallels the magnetic profiles. The

C ages on samples

from the core show that all the major changes that took place began

10 ka ago at the end of the last ice age. Analysis of pollens showed

a reduction in coniferous types as more deciduous types were trans-

ported into the lake accompanying the climatic warming in the early

Holocene. These changes are reflected in the magnetite content of

the lake sediments. The coercivity parameters provide magnetic strati-

graphies that reflect directly the paleoclimatic changes.

William Lowrie

Bibliography

Alvarez, W., Arthur, M.A., Fischer, A.G., Lowrie, W., Napoleone, G.,

Premoli Silva, I., and Roggenthen, W.M., 1977. Upper Cretaceous-

Paleocene magnetic stratigraphy at Gubbio, Italy. V. Type section

for the late Cretaceous-Paleocene geomagnetic reversal timescale.

Geological Society of America Bulletin, 88: 383–389.

Cande, S.C., and Kent, D.V., 1992a. A new geomagnetic polarity time-

scale for the Late Cretaceous and Cenozoic. Journal of Geophysi-

cal Research, 97: 13917–13951.

Cande, S.C., and Kent, D.V., 1992b. Ultra-high resolution marine

magnetic anomaly-profiles: a record of continuous paleointensity

variations? Journal of Geophysical Research, 97: 15075–15083.

Cande, S.C., and Kent, D.V., 1995. Revised calibration of the geomag-

netic polarity timescale for the Late Cretaceous and Cenozoic.

Journal of Geophysical Research, 100: 6093–6095.

Cande, S.C., and LaBrecque, J.L., 1974. Behaviour of the Earth’s

palaeomagnetic field from small scale marine magnetic anomalies.

Nature, 247:26– 28.

Channell, J.E.T., and Erba, E., 1992. Early Cretaceous polarity chrons

CM0 to CM11 recorded in northern Italian land sections near Bres-

cia. Earth and Planetary Science Letters, 108: 161–179.

Channell, J.E.T., Erba, E., Nakanishi, M., and Tamaki, K., 1995. Late

Jurassic-Early Cretaceous timescales and oceanic magnetic anom-

aly block models. In Berggren, W.A., Kent, D.V., Aubry, M., and

Hardenbol, J. (eds.), Geochronology, Timescales, and Global Stra-

tigraphic Correlation. SEPM Special Publication, pp. 51–64.

Channell, J.E.T., Erba, E., Muttoni, G., and Tremolada, F., 2000. Early

Cretaceous magnetic stratigraphy in the APTICORE drill core and

adjacent outcrop at Cismon (Southern Alps, Italy), and correlation

to the proposed Barremian-Aptian boundary stratotype. Geological

Society of America Bulletin, 112: 1430–1443.

Channell, J.E.T., Lowrie, W., and Medizza, F., 1979. Middle and Early

Cretaceous magnetic stratigraphy from the Cismon section, north-

ern Italy. Earth and Planetary Science Letters, 42: 153–166.

Channell, J.E.T., and Medizza, F., 1981. Upper Cretaceous and Paleo-

gene magnetic stratigraphy and biostratigraphy from the Venetian

(Southern) Alps. Earth and Planetary Science Letters, 55:

419–432.

Heller, F., and Evans, M.E., 1995. Loess magnetism. Reviews of Geo-

physics, 33: 210–240.

Heller, F., and Liu, T.S., 1984. Magnetism of Chinese loess deposits.

Geophysical Journal of the Royal Astronomical Society,

77:

125–141.

Horner, F., and Heller, F., 1983. Lower Jurassic magnetostratigraphy at

the Breggia Gorge (Ticino, Switzerland) and Alpe Turati (Como,

Italy). Geophysical Journal of the Royal Astronomical Society,

73: 705–718.

Kent, D.V., Olsen, P.E., and Witte, W.K., 1995. Late Triassic-earliest

Jurassic geomagnetic polarity sequence and paleolatitudes from

drill cores in the Newark rift basin, eastern North America. Journal

of Geophysical Research, 100: 14965–14998.

Krijgsman, W., Hilgen, F.J., Langereis, C.G., Santarelli, A., and

Zachariasse, W.J., 1995. Late Miocene magnetostratigraphy, bios-

tratigraphy and cyclostratigraphy in the Mediterranean. Earth and

Planetary Science Letters, 136: 475–494.

Lanci, L., Hirt, A.M., Lowrie, W., Lotter, A.F., Lemcke, G., and

Sturm, M., 1999. Mineral magnetic record of Late Quaternary cli-

matic changes in a high Alpine lake. Earth and Planetary Science

Letters, 170:49–59.

Lowrie, W., 1988. Magnetostratigraphy and the geomagnetic polarity

record. Cuàdernos de Geologia Ibérica , 12:95–120.

Lowrie, W., and Alvarez, W., 1977. Late Cretaceous geomagnetic

polarity sequence: detailed rock- and palaeomagnetic studies of

the Scaglia Rossa limestone at Gubbio, Italy. Geophysical Journal

of the Royal Astronomical Society, 51: 561–581.

Lowrie, W., and Alvarez, W., 1981. One hundred million years of geo-

magnetic polarity history. Geology, 9: 392–397.

Lowrie, W., and Alvarez, W., 1984. Lower Cretaceous magnetic strati-

graphy in Umbrian pelagic limestone sections. Earth and Plane-

tary Science Letters, 71: 315–328.

Lowrie, W., Alvarez, W., Premoli Silva, I., and Monechi, S., 1980.

Lower Cretaceous magnetic stratigraphy in Umbrian pelagic carbo-

nate rocks. Geophysical Journal of the Royal Astronomical Society,

60: 263–281.

Lowrie, W., Alvarez, W., Napoleone, G., Perch-Nielsen, K., Premoli

Silva, I., and Toumarkine, M., 1982. Pale ogene magnetic strati-

graphy in Umbrian pelagic carbonate rocks: the Contessa sec-

tions, G ubbio. Geological Society of America Bulletin, 93:

414–432.

Lowrie, W., and Channell, J.E.T., 1984. Magnetostratigraphy of the

Jurassic-Cretaceous boundary in the Maiolica limestone (Umbria,

Italy). Geology, 12:44–

47.

Lowrie, W., and Ogg, J.G., 1986. A magnetic polarity timescale for

Early Cretaceous and Late Jurassic. Earth and Planetary Science

Letters, 76: 341–349.

Mead, G.A., 1996. Correlation of Cenozoic-Late Cretaceous geomag-

netic polarity timescales: an Internet archive. Journal of Geophysi-

cal Research, 101: 8107–8109.

Moreau, M.-G., Bucher, H., Bodergat, A.-M., and Guex, J., 2002.

Pliensbachian magnetostratigraphy: new data from Paris Basin

(France). Earth and Planetary Science Letters, 203: 755– 767.

Ogg, J.G., 1984. Comment on: Magnetostratigraphy of the Jurassic-

Cretaceous boundary in the Maiolica limestone (Umbria, Italy).

Geology, 12: 701.

Ogg, J.G., Steiner, M.B., Oloriz, F., and Tavera, J.M., 1984. Jurassic

magnetostratigraphy. 1. Kimmeridgian-Tithonian of Sierra Gorda

and Carcabuey, southern Spain. Earth and Planetary Science

Letters, 71: 147–162.

MAGNETOSTRATIGRAPHY 669

Opdyke, N.D., 1972. Paleomagnetism of deep-sea cores. Reviews of

Geophysics and Space Physics, 10: 213 –249.

Opdyke, N.D., and Channell, J.E.T., 1996. Magnetic Stratigraphy. San

Diego, CA: Academic Press, 346 pp.

Parés, J.M., and Lanci, L., 2004. A Middle Eocene-Early Miocene mag-

netic polarity stratigraphy in Equatorial Pacific sediments (ODP

Site 1220). In Channell, J.E.T., Kent, D.V., Lowrie, W., and

Meert, J. (eds.), Timescales of the Paleomagnetic Field. American

Geophysical Union, Geophysical Monograph Series, 145, American

Geophysical Union, Washington, DC, pp. 130–140.

Prévot, M., Mankinen, E.A., Grommé, C.S., and Coe, R.S., 1985.

How the geomagnetic field vector reverses polarity. Nature, 316:

230–234.

Sager, W.W., Weiss, C.J., Tivey, M.A., and Johnson, H.P., 1998. Geo-

magnetic polarity reversal model of deep-tow profiles from the

Pacific Jurassic Quiet Zone. Journal of Geophysical Research,

103: 5269–5286.

Steiner, M.B., Ogg, J.G., Melendez, G., and Sequeiros, L., 1985.

Jurassic magnetostratigraphy. 2. Middle-Late Oxfordian of Aguilon,

Iberian Cordillera, northern Spain. Earth and Planetary Science

Letters, 76:151–166.

Steiner, M., Ogg, J., and Sandoval, J., 1987. Jurassic magnetostratigra-

phy. 3. Bathonian-Bajocian of Carcabuey, Sierra Harana and Cam-

pillo de Arenas (Subbetic Cordillera, southern Spain). Earth and

Planetary Science Letters, 82: 357–372.

Cross-references

Geomagnetic Polarity Reversals

Geomagnetic Polarity Timescales

Magnetic Susceptibility

Magnetization, Isothermal Remanent (IRM)

Magnetization, Remanent

Superchrons, Changes in Reversal Frequency

MAGNETOTELLURICS

Introduction

Magnetotellurics (MT) is the use of natural electromagnetic signals to

image subsurface electrical conductivity structure through electromag-

netic induction. The physical basis of the magnetotelluric method was

independently discovered by Tikhonov (1950) and Cagniard (1953).

After a debate over the length scale of the incident waves the techni-

que became established as an effective exploration tool (Vozoff,

1991; Simpson and Bahr, 2005).

Basic metho d of magnetotellurics

Electromagnetic waves are generated in the Earth’s atmosphere

and magnetosphere by a range of physical processes (Vozoff, 1991).

Below a frequency of 1 Hz, most of the signals originate in the mag-

netosphere as periodic external fields including magnetic storms and

substorms and micropulsations. These signals are normally incident

on the Earth’s surface. Above a frequency of 1 Hz, the majority of

electromagnetic signals originate in worldwide lightning activity.

These signals travel through the resistive atmosphere as waves

and when they strike the surface of the Earth most of the signal is

reflected. However, a small fraction is transmitted into the Earth and

is refracted vertically downward, owing to the decrease in propagation

velocity (Figure M168). The oscillating magnetic field of the wave

generates electric currents in the Earth through electromagnetic induc-

tion, and the signal propagation becomes diffusive, resulting in signal

attenuation with depth. The signals diffuse a distance into the Earth

that is defined as the skin depth, d, in meters by

d ¼

ﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃﬃ

pmsf



503

ﬃﬃﬃﬃﬃ

where s is the conductivity (S m

1

), f is the frequency (Hz), and m is

the magnetic permeability. The skin depth is inversely related to the

frequency and thus low frequencies will penetrate deeper into the

Earth. The impedance of the Earth is defined as

where E

is the horizontal electric field and H

is the orthogonal, hor-

izontal, magnetic field. From this impedance, apparent resistivity can

be defined as

2pf m

and the electric and magnetic field components will have a phase

difference

¼ argðZ

Þ:

In general, the impedance is written as a tensor,



;

which relates the horizontal components of the electric and magnetic

fields. The impedance relates the applied magnetic fields to the result-

ing electric fields and can be considered a transfer function. Note that

the apparent resistivity depends on the ratio of the electric and mag-

netic field components. This makes the MT method simple to apply

by combining values of E

and H

recorded at different times. The

other useful characteristic is that the direction of the incident wave

does not affect the value of apparent resistivity. The apparent resistiv-

ity can be considered an average value of the Earth’s resistivity over a

hemisphere of radius d. Thus, by computing apparent resistivity as

Figure M168 Propagation of electromagnetic waves from a

distant lightning strike to the location where MT data is recorded.

The resistive atmosphere forms a waveguide between the

conductive Earth and ionosphere. The electric field (E ) and

magnetic field (H) are both orthogonal to the direction of

propagation (k). Note that the electromagnetic energy travels as a

wave in the atmosphere, but diffuses in the Earth. This type of

signal propagation occurs above 1 Hz.

670 MAGNETOTELLURICS

a function of frequency, the variation of resistivity with depth can be

determined. This is illustrated in Figure M169. At high frequency

(1000–10 Hz) the apparent resistivity is equal to the true resistivity

of the upper layer. As the frequency decreases, the skin depth increases

and the MT signal penetrates further into the Earth and the apparent

resistivity rises. The MT phase (F

) is the phase delay between the

electric and magnetic fields at the Earth’s surface. The apparent resis-

tivity and phase are related through

¼ 45 1 þ

dðlog

f Þ



where F

is in degrees. Thus, when the apparent resistivity increases

with decreasing frequency, the phase will be less than 45



. Similarly,

a decrease in resistivity will correspond to a phase greater than 45



At the lowest frequency, the apparent resistivity asymptotically

approaches the true resistivity of the lower layer, and the phase returns

to 45



. Note that the phase is sensitive to changes in subsurface resis-

tivity with depth. For a multilayer model, MT data can reliably deter-

mine the conductance of a layer. Conductance is the vertically

integrated conductivity, and for a uniform layer the conductance is

the product of conductivity and thickness. A consequence of the

inverse problem of electrical conductivity is that MT data cannot indi-

vidually determine the conductivity and thickness of a layer. Thus

layers with differing values of conductivity and thickness, but the

same overall conductance cannot be distinguished with MT.

Early studies analyzed MT data in terms of a one-dimensional (1D)

conductivity model. In this class of model, conductivity only varies

with depth. This approach is sometimes valid in locations where the

geoelectric structure does not change rapidly in the horizontal

direction. However, it is usually necessary to consider at least a two-

dimensional (2D) Earth model. In this case, the apparent resistivity

computed from E

and H

will differ from that derived from E

and

and the application of a 1D MT analysis can give misleading

results. For a 2D Earth, E

is dependent only on H

and H

, and these

three field components comprise the transverse electric (TE) mode

with the impedance (Z

) computed from E

and H

. The transverse

magnetic (TM) mode comprises the H

, E

, and E

field components,

with the impedance (Z

) computed from E

and H

(Figure M170).

In a 2D Earth with the x-axis parallel to the geoelectric strike direction,

the impedance tensor can be written as:



0 Z



The TE mode is most sensitive to along-strike conductors. In the

TM mode the electric current flows across the boundaries between

regions of differing resistivities, which causes electric charges to build

up on interfaces. Thus the TM mode is more effective than the TE

mode at locating interfaces between regions of differing resistivity.

If the subsurface structure is three-dimensional (3D) then all four

elements of the impedance tensor are nonzero. Progress has been made

in the last decade in 3D MT modeling and inversion. However, if a

single profile of MT stations is available, and 3D effects can be shown

to be small, then a 2D analysis can be used. If the subsurface conduc-

tivity structure exhibits electrical anisotropy, this will influence the

measured impedance tensor. However, it can be difficult to con-

vincingly distinguish heterogeneity from anisotropy in MT data.

Small-scale, near-surface bodies can generate electric charges on their

boundaries. If the body is small, then insignificant electromagnetic

induction occurs and the only effect is galvanic distortion. This

changes the magnitude of the electric field at the surface and can cause

a static shift, which is a frequency-independent offset in the apparent

resistivity curve (Jones, 1988). The phase curve is not affected. Static

shifts are an example of spatial aliasing. A range of techniques is used

to remove static shifts, and include external measurements of surface

resistivity and estimation of the static shift coefficient in modeling

and inversion.

Magnetotelluric data collection and

time series processing

MT data are recorded in the time-domain, with the electric fields mea-

sured using dipoles 50–200 m in length that are connected to the

ground with nonpolarizing electrodes. Audiomagnetotelluric (AMT)

data (10000–1 Hz) typically sample the upper 1–2 km and are often

used in mineral exploration (see EM, industrial uses). Magnetic fields

are measured with induction coils, and in noisy environments the nat-

ural signals are supplemented with a transmitter. This modified techni-

que is termed controlled-source audio magnetotellurics (CSAMT).

Broadband MT data (1000–0.001 Hz) are used for sounding to mid-

crustal depths. Induction coils are generally used and a recording time

of one day is required. In the presence of excessive cultural noise addi-

tional recording may be needed. Noise can originate in a wide range of

Figure M169 Variation of apparent resistivity and phase that

would be measured at the surface of a two-layer Earth model.

Note that the depth sounding of resistivity is achieved by varying

the frequency of the signal. The dip in apparent resistivity below

10 Om at 1 Hz is a resonance phenomenon.

Figure M170 Geometry of electromagnetic field components

over a two-dimensional Earth. The transverse electric (TE) mode is

also called the E-polarization. Similarly, the transverse magnetic

(TM) mode is also called the B-polarization.

MAGNETOTELLURICS 671

sources, including power lines, cathodically protected pipelines, rail-

ways, water pumps, and electric fences. Long-period magnetotelluric

(LMT) data measure very low frequencies (1–0.0001 Hz) and are used

for imaging the lower crust and upper mantle. A specialized LMT

instrument is used with a fluxgate magnetometer, solar panels, and

low power electronics. MT data can also be collected on the seafloor.

MT time series data are processed to yield frequency-domain esti-

mates of apparent resistivity and phase. Modern processing schemes

compute fast Fourier transforms of subsections of the time series and

then utilize robust statistical techniques to average the multiple esti-

mates of the impedance. The application of robust statistics has drama-

tically improved the quality of responses and allowed many types of

noise to be effectively suppressed (Jones et al. 1989; Egbert, 1997).

In MT data collection, time series data should be recorded simulta-

neously at several locations to allow for the removal of noise at the

measurement location through the remote-reference method (Gamble

et al. 1979). This is important even in locations with minimal cultural

noise (Figure M171). In this example, ground motion from ocean waves

caused oscillations of the magnetic sensors and resulted in the apparent

resistivity being artificially low in the band 3–0.3 Hz. When the data

were processed with a remote reference, the bias was removed.

Magnetotelluric data interpret ation

Before modeling or inverting MT data, it is vital to understand the

dimensionality of MT data. Tensor decomposition is a common

approach and several techniques exist (Groom and Bailey, 1989; Bahr,

1991). Each method determines how well the measured MT impe-

dance data can be fit to a 2D geoelectric model and gives an estimate

of the geoelectric strike direction. It is common for a well-defined,

consistent geoelectric strike to only be defined in a subset of an MT

dataset. Decomposition can also determine if shallow conductivity

structures are causing galvanic distortion of the surface electric fields.

Galvanic distortion is a more general case of the process responsible

for static shifts, and changes both the magnitude and direction of the

electric fields. In extreme cases, near-surface structure can also cause

distortion of the magnetic field through intense current channeling

(Lezaeta and Haak, 2003).

Once the dimensionality has been understood, and distortion

addressed, MT can be forward modeled or inverted in 1D, 2D, or 3D

to recover a model of subsurface electrical conductivity. The inverse

problem of electrical conductivity is nonunique (Berdichevsky and

Dmitriev, 2002), which implies that a finite set of MT data containing

noise can be reproduced by an infinite number of geoelectric models.

To overcome this nonuniqueness and select a preferred model, addi-

tional constraints must be imposed on the solution. One of the most

successful methods is to require that the geoelectric model derived

from the inversion satisfies both the MT data and some additional

requirements (regularization). In the absence of any other geoelectric

information, the most common requirement is that the resistivity

model should be as spatially smooth as possible in the horizontal

and vertical directions (Constable et al., 1987). Widely used inversion

algorithms for MT data include those of Rodi and Mackie (2001) and

Siripunvaraporn and Egbert (2000). The fit of the model is usually

measured in term of the root-mean-square (rms) misfit of the predicted

model response to the measured data. An rms misfit significantly

greater than one indicates that the inversion is incapable of fitting

the MT data, and usually indicates excessive noise in the data, or 3D

effects that cannot be physically reproduced by a 2D inversion algo-

rithm. A misfit significantly less than one indicates that either the error

bars were too large or that the data is being over fit. In this second sce-

nario, the resistivity model usually appears spatially rough, with the

appearance of a checkerboard.

The MT method is now routinely used in both commercial explora-

tion and in research. Commercial applications include exploration for

minerals, hydrocarbons, and geothermal resources (see EM, industrial

uses). Researchers use MT to study the structure of the continents and

the dynamics of plate boundaries (Brown, 1994) and also in EM,

regional studies. MT measurements are also made on the seafloor

for both commercial and academic investigations (see EM, marine

controlled source).

Martyn Unsworth

Bibliography

Bahr, K., 1991. Geological noise in magnetotelluric data: a classifica-

tion of distortion types. Physics of the Earth and Planetary Inter-

iors, 66:24–38.

Berdichevsky, M.N., and Dmitriev, V.I., 2002. Magnetotellurics in the

Context of the Theory of Ill-posed Problems. Tulsa, OK: Society of

Exploration Geophysicists.

Brown, C., 1994. Tectonic interpretation of regional conductivity

anomalies. Surveys in Geophysics, 15: 123–157.

Cagniard, L., 1953. Basic theory of the magnetotelluric method of

geophysical prospecting. Geophysics, 18 : 605– 635.

Figure M171 Estimates of apparent resistivity and phase at an MT station in central California in 1999. The open circles were derived

from remote reference processing, while the black circles were derived from local data only. The MT data are contaminated by

magnetic noise due to ocean wave-induced ground motion. Note the downward bias in the apparent resistivity in the frequency band

3–0.3 Hz when local MT data processing is used.

672 MAGNETOTELLURICS

Constable, S.C., Parker, R.L., and Constable, C.G., 1987. Occam’s

inversion: a practical algorithm for generating smooth models from

electromagnetic sounding data. Geophysics, 52: 289–300.

Egbert, G.D., 1997. Robust multiple-station magnetotelluric data pro-

cessing. Geophysical Journal International, 130: 475–496.

Gamble, T.B., Goubau, W.M., and Clarke, J., 1979. Magnetotellurics

with a remote reference. Geophysics, 44:53–68.

Groom, R.W., and Bailey, R.C., 1989. Decomposition of magneto-

telluric impedance tensors in the presence of local three-dimensional

galvanic distortion. Journal of Geophysical Research, 94:1913–1925.

Jones, A.G., 1988. Static shift of MT data and its removal in a sedi-

mentary basin environment. Geophysics, 53: 967–978.

Jones, A.G., Chave, A.D., Egbert, G.D., Auld, D., and Bahr, K., 1989.

A comparison of techniques for magnetotelluric response function

estimates. Journal of Geophysical Research, 94: 14201–14213.

Lezaeta, P., and Haak, V., 2003. Beyond magnetotelluric decomposi-

tion: induction, current channeling, and magnetotelluric phases

over 90



. Journal of Geophysical Research, 108, doi:10.1029/

2001JB000990.

Rodi, W., and Mackie, R.L., 2001, Nonlinear conjugate gradients algo-

rithm for 2-D magnetotelluric inversion. Geophysics, 66: 174.

Simpson, F., and Bahr, K., 2005. Practical Magnetotellurics. Cam-

bridge: Cambridge University Press, p. 270.

Siripunvaraporn, W., and Egbert, G.D., 2000. An efficient data-

subspace inversion for two-dimensional magnetotelluric data.

Geophysics, 65: 791–803.

Tikhonov, A.N., 1950. Determination of the electrical characteristics

of the deep strata of the Earth’s crust. Doklady Akademii Nauk,

SSSR, 73(2): 295–297.

Vozoff, K., 1991. The Magnetotelluric method. In Nabighian, M.N.

(ed.), Electromagnetic Methods in Applied Geophysics, Vol. 2,

Chapter 8. Tulsa, OK: Society of Exploration Geophysicists.

Cross-references

Anisotropy, Electrical

EM Modeling, Forward

EM Modeling, Inverse

EM, Industrial Uses

EM, Marine Controlled Source

EM, Regional Studies

Galvanic Distortion

Magnetometers, Laboratory

Periodic External Fields

Storms and Substorms, Magnetic

Transfer Functions

MAGSAT

NASA (US) scientific satellite (1979–1980) made the first precise,

globally distributed measurements of the vector magnetic field near

Earth. The satellite flew at an altitude of 300–550 km, in a near-polar

inclination. Researchers modeling the Magsat data have developed

representations of the internal field extending from spherical harmonic

(q.v.)1–36 for the vector field (Lowe et al., 2001), with additional

resolution to 65 for the scalar field (Sabaka et al., 2004). Magsat also

contributed significantly to the understanding of the ionospheric mag-

netic field, and the meridional current systems through which the satel-

lite flew. The most demanding aspect of making vector magnetic

measurements from space is in precise alignment, which must be at

the arc-second level. Magnetic field measurements have been made

from near-Earth space since 1958, when Sputnik 3 (USSR) collected

fluxgate magnetometer measurements of the field with accuracies of

about 100 nT. Magsat’s measurements of the near-Earth vector field

were not improved upon, or even repeated, until the launchings of

Ørsted (q.v.) (1999) and CHAMP (q.v.) (2000), led by Denmark and

Germany, respectively. Follow-on magnetic field missions by NASA

focused on planetary magnetic fields (e.g., Mercury, Venus, Mars,

and the outer planets), on the Earth’s magnetosphere, and the

dynamics of the high-latitude ionosphere.

The Magsat project scientist was Robert A. Langel (q.v.), the project

manager was Gilbert Ousley, and the spacecraft design, development,

and testing was done by the Applied Physics Laboratory (APL) of

John Hopkins University (Potemra et al., 1981). The vector magnet-

ometer was designed and built at Goddard Space Flight Center’s

Laboratory for Extraterrestrial Physics by Mario Acuña (instrument

scientist) and his group. Scientists and scientific organizations in the

United States, Europe, and Japan formed the largest group of users

of Magsat observations. Magsat weighed 182 kg, and cost 19.2 million

US dollars from inception through production of the finished data

products. Magsat was modeled on the Small Astronomical Satellite

(SAS-3), built by APL in 1975. SAS-3 utilized a Doppler tracking

system for position determination to within 60 m vertically and 200 m

horizontally, and had two star trackers that could provide attitude infor-

mation to 10

. However, the magnetic fields associated with these star

trackers and the spacecraft itself would have introduced unacceptably

large errors into the magnetic field determinations. Therefore, a deploy-

able scissors boom and an infrared attitude transfer system (ATS) were

developed for Magsat that put the scalar and vector magnetometers six

meters away from the star trackers, and the body of the spacecraft

(Figure M172). The deployable boom was designed, developed, and

tested by APL, and the ATS was based on one used for submarines

and adapted for Magsat by the Barnes Engineering Company. The sun-

synchronous orbit of Magsat, resulting in a sampling of the magnetic

field only at dawn and dusk local time, was a compromise dictated by

the carrying capacity of the Scout launch vehicle. The chosen orbit,

and mission lifetime (October 1979 through June 1980) allowed for a

maximum exposure of the fixed solar array to the sun, a near-constant

thermal environment, and a fixed flight attitude that allowed the star

trackers to always face away from the earth.

The orientation of the vector magnetometer was determined, at 0.25 s

intervals, using the two star cameras. If only one star camera was track-

ing an identified star, a precision sun sensor immediately adjacent to

the vector magnetometer was used instead. Jumps in the vector data of

10–15 nT can occur at locations where the method of determining the

orientation of the vector magnetometer changes. The system for transfer-

ring the attitude determined by the star camera on the spacecraft body to

the vector magnetometer on the end of the boom involved two optical

systems, one for pitch and yaw (utilizing a plane mirror), and a second

for roll (using a dihedral mirror). As predicted, the system for measuring

roll had the largest errors, estimated at about eight times the pitch and

yaw error.

Magsat carried both scalar and vector magnetometers. The scalar

magnetometer, a two-sensor, cesium 133 vapor optically pumped mag-

netometer, was used to calibrate the vector magnetometer. The two-

sensor arrangement was chosen to minimize null zones, where the

magnetometers are incapable of sensing the ambient field. In practice,

intermittent noise in the cesium lamp excitation circuits caused

the tracking filters to lose lock, and reduced the amount of data

returned to well below the nominal 8 Hz rate. System errors in the

determination of the scalar field were 0.5–1.5 nT. The vector instru-

ment was a triaxial fluxgate magnetometer with a dynamic range of

2000 nT, and digitally controlled current sources to increase its range

to 64000 nT. The vector instrument sampled the ambient field

16 times per second with a resolution of 0.5 nT. Each of the three

orthogonal sensors was wound with platinum wire on a highly perme-

able toroidal magnetic core. The sensors were mounted on a stable

ceramic block and temperature controlled to minimize alignment shifts

with respect to the reference mirrors of the ATS.

Each observation has associated with it an attitude uncertainty, and

details on the sensors that were used to calculate the attitude (Langel

MAGSAT 673

et al., 1981). The mission requirements specified that the three axes of

the vector magnetometer be determined within 60

, although in prac-

tice it was often known within 20

, corresponding to an attitude error

of 5 nT. A detailed error budget can be found in Langel and Hinze

(1998) in Table 3.4. The observations can be accessed (as daily files)

at the FTP server (ftp.spacecenter.dk/data/magnetic-satellites/Magsat)

of the Danish National Space Center.

Michael E. Purucker

Bibliography

Langel, R.A., Berbert, J., Jennings, T., and Horner, R., 1981. Magsat

Data Processing: A report for investigators, NASA Technical

Memorandum 82160.

Langel, R.A., and Hinze, W.J., 1998. The Magnetic Field of the Earth’s

Lithosphere: The Satellite Perspective. Cambridge: Cambridge

University Press.

Lowe, D.A.J., Parker, R.L., Purucker, M.E., and Constable, C.G.,

2001. Estimating the crustal power spectrum from vector Magsat

data. Journal of Geophysical Research, 106: 8589–8598.

Potemra, T.A. et al., 1980. John Hopkins APL Technical Digest. The

Magsat issue, July-September, 1980. 12 articles, pp. 162– 232.

Sabaka, T.J., Olsen, N., and Purucker, M., 2004. Extending compre-

hensive models of the Earth’s magnetic field with Oersted and

CHAMP data. Geophysical Journal International, 159: 521–547.

Cross-references

CHAMP

Harmonics, Spherical

Langel, Robert A. (1937–2000)

Laplace’s Equation, Uniqueness of Solutions

Ørsted

POGO (OGO-2, -4, and -6 Spacecraft)

MAIN FIELD MAPS

Historically, in navigating long distances, the compass was as impor-

tant an instrument for indicating direction as the sandglass was for

marking time. However, until the 17th century, everyone believed that

magnetic north coincided with true north, and mariners were unaware

of this variation (see Geomagnetism, history of ); consequently, as they

sailed west they found their position did not correspond with their

location according to charts. Having learned that true north differed

from magnetic north, instrument makers in some of the northern

countries produced compasses in which the compass card was

mounted on the magnetic needle in alignment with the amount of mag-

netic variation. Thus, while the needle pointed to magnetic north, the

fleur-de-lis on the compass card indicated true north. By this arrange-

ment the compass had a built-in correction for the magnetic variation.

This was fine, as long as voyages were limited to regions where the

amount of variation did not appreciably change.

The earliest magnetic charts—before 1700

As the geomagnetic poles lie at some distance from the geographical

poles, the deviation of the compass needle (known as the “magnetic

variation” by mariners and “magnetic declination” by geophysicists)

can vary considerably over the Earth’s surface and can be directed

either to the East (positive) or to the West (negative).

What led people to include magnetic information on charts? The

first practical reason was navigation, and the earliest map that included

magnetic information was a road map, around 1500, indicating the pil-

grim’s route from Germany to Rome (Das ist der Rom Weg). On this

map, Erhard Etzlaub represented a pictorial pocket sundial (known at

that time as a compass), on which he clearly showed an easterly decli-

nation (Hellmann, 1909).

During the 16th century, the number of voyages increased, and nau-

tical charts were being drawn. In 1576, William Borough produced a

chart for Martin Frobisher’s first voyage in search of a North-West

passage. On this manuscript, numerical values for declination were indi-

cated for several locations situated along the likely trip (Barraclough,

2000). A global map indicating magnetic poles, magnetic meridians,

and the magnetic equator was published by Guillaume de Nautonnier

(1602–1604). In Edward Wright’s book (1610), a chart with values

for declination is included. Between 1625 and 1634, Jean Guérard, a

French pilot and hydrographer, produced seven world maps; one of

them includes declination values, marked in small circles at the places

of measurements (Ultré-Guérard and Mandea, 2000).

If knowledge of the declination was acquired during the 15th cen-

tury, the variation of declination with geographical position became

more widely known around the end of 15th century. One very interest-

ing application was to use the declination to find the longitude at sea.

A sailor can find the latitude by the length of day or by the height of

the Sun or other known star above the horizon. To measure longitude

one needs to know, at the same moment, the time at the departure

place, and the time aboard ship. Then, measuring longitude means

measuring the time. However, until John Harrison’s (1760) very pre-

cise ship-chronometer, it was not obvious how to measure time with

enough accuracy. One proposed method was to use the magnetic

compass, and thus declination, to find the longitude. According to

Hellman (1909) Columbus was the first to have the idea, but, this

has since been discredited (Mitchell, 1937). Mitchell indicated that in

1544, Sebastian Cabot (Pilot Major of Spain) produced a planisphere

on which the contour of zero declination was indicated. Another

Figure M172 Line drawing of Magsat, its orientation, and critical systems. The roll axis is along the orbital velocity vector, the yaw axis

is vertical, and the pitch axis completes the orthogonal coordinate system.

674 MAIN FIELD MAPS

attempt to solve the longitude problem using declination was made by

Guillaume de Nautonnier (1602–1604); for more details about Nau-

tonnier’s map and work see Mandea (2000). In another book, pub-

lished in 1609 by Antony Linton (strongly influenced by

Nautonnier), it was suggested that the declination resulting from a

model with different positions for the magnetic and geographic poles

could be represented on a sphere by “helique and tortuos lines”.

The earliest magnetic charts—1700 onward

Feeling the need for more accurate magnetic charts of the Atlantic

Ocean, the lordships of the British Admiralty lent Edmund Halley a

small sailing ship, Paramore, and instructed him to carry out a magnetic

survey of the Atlantic Ocean and its bordering lands. From Halley’s

survey measurements the first magnetic map of the Atlantic was com-

piled (Figure M173/Plate 1). This earliest surveying, magnetic contour

Figure M173/Plate 1 Halley’s Atlantic chart. This is the version of the chart with the cartouche over Africa containing a dedication to

King William III.

MAIN FIELD MAPS 675