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

Подождите немного. Документ загружается.

Dean, W.C., 1958. Frequency analysis for gravity and magnetic inter-

pretation. Geophysics, 23:97–127.

Emilia, D.A., 1973. Equivalent sources used as an analytic base for

processing total magnetic field profiles. Geophysics, 38: 339–348.

Fedi, M., Rapolla, A., and Russo, G., 1999. Upward continuation of

scattered potential field data. Geophysics, 64: 443–451.

Henderson, R.G., 1970. On the validity of the use of the upward con-

tinuation integral for total magnetic intensity data. Geophysics, 35:

916–919.

Henderson, R.G., and Zietz, I., 1949. The upward continuation of

anomalies in total magnetic intensity fields. Geophysics, 14:517–534.

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

Earth’s Lithosphere: The Satellite Perspective. Cambridge:

Cambridge University Press.

Pawlowski, R.S., 1995. Preferential continuation for potential-field

anomaly enhancement. Geophysics, 60: 390–398.

Ravat, D., Whaler, K.A., Pilkington, M., Purucker, M., and Sabaka, T.,

2002. Compatibility of high-altitude aeromagnetic and satellite-

altitude magnetic anomalies over Canada. Geophysics, 67:546–554.

Shure, L., Parker, R.L., and Backus, G.E., 1982. Harmonic splines for

geomagnetic modeling. Physics of the Earth and Planetary Inter-

iors, 28: 215–229.

Cross-references

Aeromagnetic Surveying

Crustal Magnetic Field

976 UPWARD AND DOWNWARD CONTINUATION

VARIABLE FIELD TRANSLATION BALANCE

The variable field translation balance (VFTB) is an instrument for

measuring isothermal magnetizations in variable fields (e.g., hysteresis

loops) as well as the temperature dependence of the associated

magnetic parameters. It is specifically designed to measure the weak

magnetizations commonly encountered in rock magnetism.

Measurement principles and main

instrument components

The VFTB is a modification of the horizontal magnetic translation bal-

ance (e.g., Collinson, 1983). In contrast to the normal horizontal trans-

lation balance, the magnetic gradient in the VFTB is not produced by a

special shape of the pole pieces of the electromagnet, but by a separate

set of four gradient coils (AC; Figure V1). This separation of inducing

field and gradient field was for the first time proposed by McKeehan

(1934). The gradient field of the VFTB is not kept constant, but is

oscillating with a certain frequency (2–4Hz). The VFTB can thus be

considered a one-dimensional harmonic oscillator with damping (D),

operated in forced oscillation mode. The oscillating part of the instru-

ment is a pendulum with bifilar suspension (BS) with the specimen (S)

fixed to it in a nonmagnetic quartz glass sample holder (Q). The motion

of the system is excited by a periodic force acting on the sample and gen-

erated by the set of gradient coils. The system is normally operated at

resonance frequency to obtain highest sensitivity, but can also be oper-

ated in a nonresonance mode for the measurement of very strong mag-

netic samples. The motion of the sample is monitored by a linear

variable differential transformer (LVDT). This signal—after processing

in the PC—is then proportional to the magnetic moment of the sample.

The temperature is controlled by a heating/cooling unit (H/C) sur-

rounding the sample chamber, thus making possible measurements at

variable temperatures. During measurement the specimen is located

in the center of either an electromagnet (EM, with pole pieces) or an

iron-free solenoid generating a homogeneous inducing magnetic field.

In case of the pole piece configuration, two compensation coils (CCs)

are attached to compensate the residual field generated by the rema-

nent magnetization of the pole pieces in the case of zero field measure-

ments. All instrument functions like inducing field strength, gradient

field amplitude, and temperature are computer controlled.

As the sample is moving back and forth inside the gradient field

its motion is detected by the LVDT vibration sensor (Figure V2).

The LVDT generates a voltage proportional to the position of the

sample holder inside the sensor. This voltage is amplified and read

by the A/D converter. The VFTB control software records the digitized

voltages from the converter board. The gradient field actuating the

sample movement varies sinusoidally. The response movement of the

sample is also shaped like a sine wave with the same frequency as

the original signal but with differing phase and amplitude. These indi-

cate the actual magnetization of the sample. To compute amplitude and

phase of the signal the recorded voltages are analyzed via fast Fourier

transformation to find the spectral amplitude at the measurement

frequency. The phase is used to compute the sign of the gained ampli-

tude thus yielding the desired magnetization value. This value is

normalized by the input gain of the A/D converter and the strength

of the actuating field.

Types of measurements

VFTB measurements can be subdivided into field-dependent, tempera-

ture-dependent, and time-dependent magnetization measurements.

The first group consists of isothermal magnetization measurements,

i.e., IRM acquisition curves, hysteresis loops, and backfield curves.

The number and spacing of field steps for these measurements can

be freely chosen depending on the sample material and the desired

measurement speed. It takes about 15min to measure a hysteresis loop

defined by 80 data points. Examples are shown in Figure V3.

The main purpose of temperature-dependent experiments is the

determination of crystallographical or physical phase transitions,

e.g., Verwey transition and Curie temperatures. The temperature can

be varied between 190



C and room temperature with liquid nitrogen

as furnace coolant and between room temperature and 800



C with

Figure V1 Schematic drawing of the main instrument components.

Drawing shows the configuration with an electromagnet with pole

pieces (see text for explanation of acronyms).

Figure V2 Simplified connection scheme of the VFTB.

FigureV3 Example data of a basaltic specimen measured with a VFTB.

978 VARIABLE FIELD TRANSLATION BALANCE

water cooling. IRM acquisition curves as well as hysteresis loops and

backfield curves can also be measured in the whole temperature range

between 190 and 800



C. The VFTB can also measure the time

dependence of magnetization at given temperatures, useful, e.g., for

determining the decay of isothermal remanent magnetization or for

observing alteration processes at elevated temperatures.

Magnetic moment sensitivity of the VFTB amounts to 5  10

8

Sample holders can accommodate samples with a diameter of either 5 or

10 mm and a length of 9mm. It is thus possible to measure fairly large

samples with a maximum weight of 1.5 g which yields a sensitivity

of  5  10

5

1

in terms of specific magnetization. Depending

on the electromagnet a maximum magnetic field of 1.0 T can be attained

and the residual field during zero field measurements does not exceed

50 m T.

David Krása, Klaus Petersen, and Nikolai Petersen

Bibliography

Collinson, D.W., 1983. Methods in Rock Magnetism and Palaeomagnet-

ism: Techniques and Instrumentation. New York: Chapman & Hall.

McKeehan, L., 1934. Pendulum magnetometer for crystal ferromag-

netism. The Review of Scientific Instruments, 5: 265– 268.

Cross-references

Magnetization, Isothermal Remanent

Rock Magnetism

Rock Magnetism, Hysteresis Measurement

VERHOOGEN, JOHN (1912–1993)

John Verhoogen (Figure V4) was an engineer, geologist, geophysicist,

and a pioneering advocate of the role of the solid inner core in the geo-

dynamo. He was also an inspirational mentor to a generation of stu-

dents, including the Earth Scientists who first established the

timescale for magnetic polarity reversals.

Verhoogen was born in Brussels in 1912 in an educated, profes-

sional family. He became afflicted with poliomyelitis while in high

school. Although this disease would progressively limit his physical

mobility, it did not deter him from pursuing a career in the geological

sciences. He earned degrees in mining engineering at the University of

Brussels in 1933, engineering geology at the University of Liege in

1934, and a Ph.D. in geology at Stanford University in 1936. This

was his first introduction to the United States, where he would later

make his scientific mark. After a decade of fieldwork in equatorial

Africa, he joined the faculty of the University of California, Berkeley

in 1947 and taught there until his retirement in 1976.

Verhoogen’s first major research accomplishment was to measure

the gas composition of the active volcano Nyamuragira in Zaire (then

the Belgian Congo). This experience convinced him that igneous pro-

cesses and volcanism were indisputable evidence for large-scale flow

deep in Earth’s interior. One outcome of these studies proposed an expla-

nation for the thermal evolution of the Earth in terms of heat transfer

by fluid motions in the deep interior. Although the significance of this

idea was not widely appreciated by his contemporaries at the time, it

guided his own research for the ensuing three decades and, ultimately,

it has to be accepted throughout the geoscience community.

In the late 1950s and early 1960s, Verhoogen played a seminal role

in the development of paleomagnetism, particularly the paleomagnetic

reversal time scale. At that time there were many reports of rocks with

ferromagnetic directions reversed to the present-day geomagnetic field

direction. Two competing explanations were offered to explain this

puzzling phenomenon: either these rocks were magnetized during

epochs when the main geomagnetic field was reversed, or their original

magnetic directions had been reversed by local remagnetization effects.

The existing techniques in geochronology were inadequate to distinguish

between the two competing explanations. Verhoogen initially favored

remagnetization, but several of his own graduate students, including

Allan Cox and Richard Doell, were measuring the paleomagnetic direc-

tions in lavas from Hawaii and other volcanic hotspots and convinced

him that the global reversal explanation was far more likely. Within a

few years, new developments in mass spectrometer technology provided

accurate radiometric dating methods, confirming that the main geo-

magnetic field had indeed reversed its polarity many times in the past.

The resulting geomagnetic polarity time scale is now a cornerstone

in geophysics; for example, it provides the interpretation of magnetic

stripes in the ocean crust in terms of sea-floor spreading (see Glen 1982

for a description of Verhoogen’s role in these developments).

As post-polio injuries curtailed his field and laboratory work,

Verhoogen progressively turned his attention toward the energy bal-

ance of the Earth as a whole, and in particular, the energy needed to

maintain the geodynamo. One of his major contributions to igneous

geology had been a quantitative thermodynamic treatment of igneous

petrogenesis, which combined the energetics of crystallization with

the thermal environment of the deep crust and upper mantle where

the magmas originate. He recognized that the core was like a giant

magma chamber, in the sense that physical and chemical processes

similar to those in magmatic systems must be operating inside the core.

This represented a new perspective; previously, only thermal convection

had been considered important in core energetics. Verhoogen reasoned

that, because the Earth as a whole is cooling, the solid inner core must

be growing by crystallization at the expense of the fluid outer core. He

realized that crystallization at the inner core boundary would provide a

source of energy for the geodynamo, through latent heat release and

other effects associated with the solidification. In 1961, he estimated

the energy release by this process, and found it adequate to maintain

the geodynamo if the rate of inner core solidification is sufficiently high,

about 25 m

1

or more. Subsequent investigations by Braginsky

(1963), Gubbins (1977), Buffett et al. (1992), and others have developed

these concepts further, adding an important ingredient to the model, the

gravitational potential energy supplied by chemical differentiation at the

crystallizing inner-outer core boundary. This model has come to be

Figure V4 John Verhoogen (courtesy of Margaret Gennaro,

University of California Archives).

VERHOOGEN, JOHN (1912–1993) 979

called the “gravitational dynamo” and is now the favored mechanism for

powering the geodynamo.

In his 1980 monograph Energetics of the Earth , Verhoogen system-

atically analyzed the various energy sources available to drive

the geodynamo, including inner core solidification, heat loss to

the mantle, and possible radioactive heat sources in the outer core.

He concluded that, although the gravitational dynamo mechanism evi-

dently provides enough energy, the uncertainties in the rate of heat loss

to the mantle, combined with the uncertainties in radioactive abun-

dance and important physical properties of the core, preclude us from

making a definitive model of the geodynamo. Researchers today

would, no doubt, come to a similar conclusion. However, in an address

to the US National Academy of Science, Verhoogen expressed his

scientific optimism in the following way:

Perhaps the greatest step forward since the discovery of radioac-

tivity has been the recognition of convection as the dominant

mode of heat and mass transfer in most of the mantle and core.

Here at long last is something to hang onto.

This statement is as applicable today as it was two decades ago, and

remains the paradigm for understanding the deep Earth.

Peter Olson

Bibliography

Braginsky, S.I., 1963. Structure of the F layer and reasons for convec-

tion in the earth’s core. Doklady Akademiya Nauk SSSR, 149:

1311–1314 (English Translation).

Buffett, B.A., Huppert, H.H., Lister, J.R., and Woods, A.W., 1992.

Analytical model for solidification of Earth’s core. Nature, 356:

329–331.

Glen, W., 1982. The Road to Jaramillo. Palo Alto, CA: Stanford Uni-

versity Press.

Gubbins, D., 1977. Energetics of the Earth’s core. Journal of Geophy-

sics, 43: 453–464.

Verhoogen, J., 1961. Heat balance of the Earth’s core. Geophysical

Journal of the Royal Astronomical Society, 4: 276–281.

Verhoogen, J., 1980. Energetics of the Earth. Washington, DC:

National Academy of Science.

Cross-references

Convection, Chemical

Cox, Allan V. (1926 –1987)

Geodynamo, Energy Sources

Geomagnetic Polarity Timescales

Magnetization, Thermoremanent (TRM)

VINE-MATTHEWS-MORLEY HYPOTHESIS

Definition

The Vine-Matthews-Morley (VMM) hypothesis states that, when

ocean crust forms at a midocean ridge (i.e., a spreading center), the

cooling crust becomes magnetized in the direction of Earth’s prevail-

ing magnetic field as it cools below the Curie temperature of the mag-

netic minerals (Morley and Larochelle, 1964; Vine and Matthews,

1963). Typically this Curie temperature is between 150 and 300



for titanomagnetite, the primary magnetic mineral in the upper oceanic

crust, and 580



C for pure magnetite found in the lower ocean crust.

The ocean crust moves away from the spreading center through the

process of “seafloor spreading” and in this way the magnetic signal

recorded in the newly formed crust is preserved (Figure V5). A good

analogy is that of a tape recorder where the ocean crust is the magnetic

tape and Earth’s magnetic field is the signal, which is recorded onto

the moving tape. In this way, changes in Earth’s magnetic field polarity

through time are preserved in the crustal rocks. The alternating bands of

magnetic polarity give rise to magnetic anomalies when measured at the

sea surface, the so-called “magnetic stripes” (Figure V6) (see Magnetic

surveys, marine; Magnetic anomalies, marine).

The VMM hypothesis elegantly explains the striped patterns of

magnetic anomalies measured at the sea surface without the need to

call upon petrological, thermal, or some other process to be aligned par-

allel to the spreading axis. It also links many of the other observations

about midocean ridges such as their apparent youth and thin crust. It

also explains why the ocean basins are relatively young (<200 My)

compared to continents. The magnetic anomalies document the progres-

sion in age of the ocean basins from young crust at midocean ridges to

the oldest and deepest crust near subduction zones and in doing so

strongly supports the global concept of continental drift that ocean

basins have opened and closed in the past. The VMM hypothesis forms

one of the cornerstones of plate tectonic theory and began a revolution in

thinking about how the Earth works and how it has evolved through geo-

logic time.

Historical context

Fred Vine and Drummond Matthews were part of the geophysics

group at Cambridge University in England at the time of their discov-

eries under the leadership of Teddy Bullard and they were very recep-

tive to the ideas of American scientists Harry Hess and Robert Dietz

(Dietz, 1961; Hess, 1962) about continental drift and the possibility

of seafloor spreading in contrast to the mainstream US science com-

munity at that time (Glen, 1982). Vine began his work as a graduate

student working for Drummond Matthews and their ideas grew out

of work by Vine on magnetic profiles collected by Matthews over

the Carlsberg Ridge in the Indian Ocean. Vine found that there were

reversely magnetized seamounts indicating that the ocean crust could

record both normal and reverse polarity periods in Earth ’s magnetic

field history. A similar conclusion had been reached earlier by Girdler

and Peter (1960) who suggested the possibility of reversely magne-

tized crust in the Gulf of Aden. Vine and Matthews (1963) suggested

that this process may be more pervasive that previously thought. The

Carlsberg Ridge profiles showed positive and negative magnetic

anomalies 20 km in wavelength, and Vine and Matthews suggested

that these anomalies could be easily explained by polarity changes in

the crustal magnetization rather than petrological or thermal differ-

ences. It was this observation that allowed Vine and Matthews to make

the daring statement that perhaps 50% of the ocean floor was reversely

magnetized. They borrowed on the ideas of Dietz (1961) and Hess

(1962) that seafloor spreading was occurring at midocean ridges and

made the key observation that ocean crust could record Earth’s mag-

netic field as it formed and that this record would then be preserved

as the crust moved away from the spreading center.

The paper initially received little fanfare, comment, or recognition.

Most scientists dismissed the paper as being too speculative. What

was not immediately appreciated was the prediction of symmetrical

anomalies about a midocean ridge crest and the fact that seafloor spread-

ing rates could be obtained from such magnetic records. Lawrence

Morley recognized these key ideas but his initial attempts at publishing

were declined. One editor reputedly commented that the paper was more

suitable as a subject for discussion at a cocktail party rather than a serious

scientific idea (Glen, 1982). Morley was an administrator for the Geolo-

gical Survey of Canada at the time and had little time to vigorously pur-

sue these ideas. Morley was eventually able to publish his article in an

obscure journal in Canada (Morley and Larochelle, 1964). None of the

significance of the VMM ideas permeated the scientific community at

the time of the publications. It was not until Vine and Wilson (1965)

looked at the newly published maps of Raff and Mason (1961)

(FigureV6) that they recognized the location of the Juan de Fuca mido-

cean ridge spreading center and its position at the center of symmetry

980 VINE-MATTHEWS-MORLEY HYPOTHESIS

in the northeast Pacific magnetic anomalies that the idea began to gain

ground and become seriously debated in academic circles. Vine and

Wilson (1965) also matched the magnetic profiles with the newly com-

piled geomagnetic polarity reversal timescale of Cox et al. (1963) based

on precisely age dated terrestrial lavas.

A subsequent paper by Vine (1966) eventually put all the ideas

together using a long magnetic profile over the East Pacific Rise (Elta-

nin Leg-19) collected by Walter Pitman of Lamont-Doherty Geologi-

cal Observatory (Pitman and Heirtzler, 1966) (Figure V7). Vine

showed that there was universal symmetry and correlation between

magnetic profiles from several midocean ridges and that these could

be modeled with the geomagnetic polarity reversal timescale of Cox

et al. (1963). This essentially sealed the case that magnetic anomalies

recorded in ocean crust preserved the history of Earth’s magnetic field

and that this reversal pattern could be correlated with dated reversals

on land.

Discoveries stemming from VMM hypothesi s

Heirtzler and colleagues (1968) took the observations of Vine and

Matthews (Glen, 1982) and Vine (1966) to the logical conclusion by

producing a geomagnetic polarity timescale (GPTS) based on the mar-

ine magnetic anomaly patterns found over the ocean basins and

extended the timescale to 80 My. Larson and Pitman (1972) and

Larson and Chase (1972) extended the GPTS with studies of Pacific

magnetic anomalies and eventually these were reconciled with anoma-

lies of a similar age in the Atlantic (Vogt et al., 1971) into the Mesozoic

or M-series anomalies. The present incarnation of the GPTS was pub-

lished by Cande and Kent (1995) with extension to the Mesozoic anoma-

lies by Channell et al. (1995). The marine magnetic record allows for a

detailed history of ocean basin formation and continental drift to be

documented in terms of plate tectonics. A detailed history of plate

motion and spreading rates has been compiled by various researchers

primarily based on the marine magnetic anomaly patterns (DeMets

et al., 1990; Minster and Jordan, 1978; Scotese, 1997). A compilation

of ship-based mapping of the marine magnetic anomaly patterns has

allowed a digital database of seafloor age to be produced (Mueller

et al., 1996).

Figure V6 Magnetic anomaly “stripe” map of northeast Pacific

Ocean, where positive anomalies are shaded black and reverse

polarity is shown in white. It was the symmetrical pattern of

anomalies about the newly discovered Juan de Fuca Ridge

spreading center that began the acceptance of the VMM

hypothesis. Figure modified from Raff and Mason (1961).

Reproduced by permission of the Geological Society of America.

Figure V5 Cartoon of the Vine-Matthews-Morley (VMM) hypothesis showing the process of seafloor spreading at a midocean ridge

and the recording Earth’s magnetic field to create the magnetic stripes. Reproduced by permission of Woods Hole Oceanographic

Institution (Tivey, 2004).

VINE-MATTHEWS-MORLEY HYPOTHESIS 981

The shape of magnetic anomalies (i.e., their phase) contains infor-

mation about their latitude of formation which can be used to deter-

mine the paleolatitude when the crust formed (Acton and Gordon,

1991; Petronotis et al., 1994). Magnetic anomaly patterns have also

led to the discovery of new types of seafloor spreading in the form

of propagating rifts (Hey, 1977; Hey et al., 1980) and rotating micro-

plates (Courtillot, 1980; Hey, 1985; Schouten et al., 1993).

Marine magnetic surveys using deep-towed sensors and autono-

mous underwater vehicles are now providing increased resolution in

mapping magnetic anomalies, which has allowed for short-wavelength

magnetic anomalies (<10 km wavelength) to be recognized (Tivey

et al., 1997). It now is apparent that the ocean crust not only records

the polarity of Earth’s magnetic field, but also provides a record of

Earth’s field intensity through time (Bowers et al., 2001; Pouliquen

et al., 2001). Magnetic profiles over ridge crests and show a good cor-

relation of short-wavelength magnetic anomalies (<10 km wave-

length) with terrestrial and sediment core derived paleointensity

records (Gee et al., 2000; Guyodo and Valet, 1999). It may, therefore,

be possible to use the paleointensity record as a fine-scale signal with

which to date the seafloor in addition to the polarity reversal signal.

Earth’s magnetic field intensity apparently undergoes relatively rapid

changes compared to the polarity reversal signal and thus the crustal

accretion process will ultimately control the resolution of the recorded

signal (Schouten et al., 1999).

Some outstanding issues

One of the outstanding and unsolved issues of the VMM hypothesis is

the source of the magnetic anomaly stripes. What part of the ocean

crust is responsible for the generating the magnetic anomalies? Vine

and Matthews (1963) made only vague stipulations as to the possible

source of the magnetic anomalies. For a crustal unit to be considered

as a viable source of the magnetic anomalies, it must satisfy a number

of criteria (Harrison, 1981; Johnson, 1979). First, the crust has to be

sufficiently magnetic in order to create the anomalies. Second, the

magnetization has to be stable and robust enough to withstand millions

of years of alteration without degrading. Third, the crust has to cool

quickly in order to capture magnetic reversals, which occur at an aver-

age frequency of 350000 years. Finally, the geometry of the reversal

boundaries must be sufficiently vertical in order to create magnetic

anomalies.

Ocean crust can be crudely thought of being composed of three

layers (Figure V8). The upper extrusive lava layer consists of lava that

has erupted onto the seafloor and thus has cooled rapidly. This layer is

highly magnetic although relatively thin, typically <1 km thick. This

layer overlies the intrusive dike layer, which are the subsurface volca-

nic conduits that feed the overlying lava flows. The dikes cool much

more slowly than the lavas and are markedly less magnetic because

of their intrinsically larger and less magnetic grains and high tempera-

ture alteration associated with their emplacement. The dike layer is on

the order of 1–2 km thick. The third crustal layer is the plutonic gab-

bro section, which can be thought of as the cooling magma chamber

that is the magmatic source of the overlying dikes and lavas. Gabbros

cool relatively slowly and have large crystals and would therefore be

intrinsically less magnetic. Nevertheless, some gabbros can be quite

magnetic with microscopic grains of pure magnetite found within pla-

gioclase crystals. The gabbros as a layer are quite thick, up to 2–4km

thick and thus while perhaps less magnetic than the extrusive lavas

could be of sufficient volume to generate a magnetic anomaly. Finally,

underlying the ocean crust is the mantle, which is composed of perido-

tite and may also be magnetic, especially if it has undergone alteration

to serpentinite through the introduction of water. The contribution of

these crustal layers and mantle to the overlying magnetic anomalies

is difficult to quantify, especially with in situ measurements. Early

magnetic surveys determined that the upper extrusive lavas were the

primary source layer based on topographic arguments (Atwater and

Figure V7 Eltanin-19 magnetic anomaly profile from Pitman and Heirtzler (1966) shown in the middle with the profile reversed

shown at top and a computed model profile shown below. The geomagnetic polarity timescale is shown at the bottom with the normal

polarity shown in black and the reverse polarity in white.

982 VINE-MATTHEWS-MORLEY HYPOTHESIS

Mudie, 1973; Talwani et al., 1971). Oceanic crustal drilling provided

ambiguous results showing that the magnetization of the extrusive

lavas was on average not enough to generate the anomalies and that

deeper layers must be contributing as well (Harrison, 1981). While

some gabbros have been drilled and are magnetic (Pariso and Johnson,

1993), ocean crustal drilling to date has not penetrated a complete sec-

tion of crust from extrusive lava to gabbro so that a complete answer

to this question remains.

The dip of a polarity boundary can also reveal useful information

about the source and contribution of the crustal layers to the magnetic

anomaly. A conceptual model of the polarity boundary (Kidd, 1977)

suggests that within the extrusive lava layer the progressive overlap

of lavas, as they are formed and are rafted away from the spreading

center, generates a boundary that dips toward the spreading center. In

the dike layer the boundary is likely to be near vertical and relatively

sharp, while in the gabbro section a polarity boundary is likely to be

a cooling isotherm, which would dip away from the spreading center

(Figure V8). A couple of detailed near-bottom magnetic surveys

showed the importance of the dip of the polarity boundary within the

extrusive layer. Macdonald et al. (1983) used the submersible ALVIN

to measure the magnetic polarity of seafloor outcrops of lava at a

reversal boundary and found a sharply defined polarity boundary

between normal and reversal magnetized lavas. This boundary location

was compared to that obtained from a towed near-bottom magnetic

survey that is an average of the entire layer contribution. The deep-

tow boundary was a kilometer closer to the spreading axis than the sur-

ficial boundary, indicating that the polarity boundary dips toward

the spreading center. A submarine magnetic survey carried out at a

fracture zone wall (Tivey, 1996; Tivey et al., 1998) mapped the mag-

netic boundaries within the extrusive layer in cross-section and also

found that the polarity boundaries dipped toward the spreading center.

In this case, the authors also were able to calculate the contribution of

the layer to the overlying anomaly and determined that the extrusive

lavas produced about 75% of the observed anomaly, indicating that

additional contribution was needed from deeper layers such as the

dikes and gabbros (Tivey et al., 1998). Only one survey has unam-

biguously determined the dip of polarity boundaries in the gabbros

to date (Allerton and Tivey, 2001). In this survey, seafloor drilling

determined the surficial polarity boundary location while a near-

bottom survey determined the average boundary position and found

that the polarity boundary dips away from the spreading center consis-

tent with the idea of cooling isotherms within the gabbro section

(Allerton and Tivey, 2001). The contribution and geometry of the

polarity boundaries within the lower ocean crust remain poorly

documented and continue to be a topic of research.

Maurice A. Tivey

Bibliography

Acton, G.D., and Gordon, R.G., 1991. A 65Ma paleomagnetic pole for

the Pacific Plate from the skewness of magnetic anomalies 27r-31.

Geophysical Journal International, 106: 407–420.

Allerton, S., and Tivey, M.A., 2001. Magnetic polarity structure of

lower oceanic crust. Geophysical Research Letters, 28: 423–426.

Atwater, T.M., and Mudie, J., 1973. Detailed near-bottom geophysical

study of the Gorda Rise. Journal of Geophysical Research, 78:

8665–8686.

Bowers, N.E., Cande, S.C., Gee, J.S., Hildebrand, J.A., and Parker, R.

L., 2001. Fluctuations of the paleomagnetic field during chron C5

as recorded in near-bottom marine magnetic anomaly data. Journal

of Geophysical Research, 106: 26379–26396.

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

netic polarity timescale for the late Cretaceous and Cenozoic. Jour-

nal of Geophysical Research, 100: 6093–6095.

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

Jurassic-Early Cretaceous time scales and oceanic magnetic anom-

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

and Hardenbol, J. (eds.), Geochronology, Time Scales and Global

Stratigraphic Correlation, Volume 54: Special Publication. Tulsa,

OK: SEPM (Society for Sedimentary Geology), pp. 51–63.

Courtillot, V., 1980. Plaques, microplaques et dechirures lithospheri-

ques: une hierarchie de structures tectoniques de l’echelle du globe

a celle du terrain. Bulletin de la Societe Geologique de France, 12:

981–984.

Cox, A., Doell, R.R., and Dalrymple, G.B., 1963. Geomagnetic

polarity epochs and Pleistocene geochronology. Nature, 198:

1049–1051.

DeMets, C., Gordon, R.G., Argus, D.F., and Stein, S., 1990. Current

plate motions. Geophysical Journal International, 101: 425–478.

Dietz, R.S., 1961. Continent and ocean basin evolution by spreading

of the seafloor. Nature, 190: 854–857.

Gee, J.S., Cande, S.C., Hildebrand, J.A., Donnelly, K., and Parker, R.L.,

2000. Geomagnetic intensity variations over the past 780kyr

obtained from near-seafloor magnetic anomalies. Nature, 408:

827–832.

Girdler, R.W., and Peter, G., 1960. An example of the importance of

natural remanent magnetization in the interpretation of magnetic

anomalies. Geophysical Prospecting

, 8: 474–483.

Glen, W., 1982. The Road to Jaramillo. Critical Years of the Revolu-

tion in Earth Science. Stanford, CA: Stanford University Press,

459 pp.

Guyodo, Y., and Valet, J.-P., 1999. Global changes in intensity of

the Earth’s magnetic field during the past 800kyr. Nature, 399:

249–252.

Figure V8 Diagram showing a cross-section through ocean crust

with the primary crustal units and thicknesses. Superimposed on

the lithology is the presumed geometry of a magnetic polarity

reversal boundary (thick black line). The boundary dips toward

the spreading axis in the extrusive lavas, but is thought to dip

away from the axis in the deeper gabbro layer following a cooling

isotherm. This portion of the geometry is hypothetical. Drawing

based on Macdonald et al. (1983) and Kidd (1977).

VINE-MATTHEWS-MORLEY HYPOTHESIS 983

Harrison, C.G.A., 1981. Magnetism of the oceanic crust. In Emiliani, C.

(ed.), The Sea, Volume 7. New York: Wiley, pp. 219–237.

Heirtzler, J.R., Dickson, G.O., Herron, E.M., Pittman, W.C., III, and

Le Pichon, X., 1968. Marine magnetic anomalies, geomagnetic

field reversals, and motions of the ocean floor and continents.

Journal of Geophysical Research, 73:2119–2136.

Hess, H.H., 1962. History of ocean basins. In Engel, A.E.J., James, H.L.,

and Leonard, B.F. (eds.) Petrologic Studies: The Buddington Volume.

Boulder, CO: Geological Society of America, pp. 599–620.

Hey, R., 1977. A new class of “pseudofaults” and their bearing on

plate tectonics: a propagating rift model. Earth and Planetary

Science Letters, 37: 321–325.

Hey, R., Duennebier, F.K., and Morgan, W.J., 1980. Propagating

rifts on midocean ridges. Journal of Geophysical Research, 85:

3647–3658.

Hey, R.N., Naar, D.F., Kleinrock, M.C., Phipps Morgan, W.J.,

Morales, E., and Schilling, J.-G., 1985. Microplate tectonics along

a superfast seafloor spreading system near Easter Island. Nature,

317: 320–325.

Johnson, H.P., 1979. Magnetization of the oceanic crust. Reviews of

Geophysics and Space Physics, 17: 215– 225.

Kidd, R.G.W., 1977. The nature and shape of the sources of marine

magnetic anomalies. Earth and Planetary Science Letters, 33:

310–320.

Larson, R.L., and Chase, C.G., 1972. Late Mesozoic evolution of the

western Pacific Ocean. Geological Society of America Bulletin,

83: 3627–3644.

Larson, R.L., and Pitman, W.C., 1972. Worldwide correlation of

Mesozoic magnetic anomalies and its implications. Geological

Society of America Bulletin, 83: 3645–3662.

Macdonald, K.C., Miller, S.P., Luyendyk, B.P., Atwater, T.M., and

Shure, L., 1983. Investigation of a Vine-Matthews magnetic linea-

tion from a submersible: the source and character of marine mag-

netic anomalies. Journal of Geophysical Research, 88: 3403–3418.

Minster, B.J., and Jordan, T.H., 1978. Present-day plate motions. Jour-

nal of Geophysical Research, 83: 5331–

5354.

Morley, L.W., and Larochelle, A., 1964. Paleomagnetism as a means

of dating geological events. Royal Society of Canada Special Pub-

lication, 8:39–50.

Mueller, R.D., Roest, W.R., Royer, J.-Y., Gahagan, L.M., and Sclater,

J.G., 1996. Age of the Ocean Floor. Report MGG-12, Data

Announcement 96-MGG-04. Boulder, CO: National Geophysical

Data Center.

Pariso, J.E., and Johnson, H.P., 1993. Do layer 3 rocks make a signif-

icant contribution to marine magnetic anomalies? In situ magneti-

zation of gabbros at ocean drilling program hole 735B. Journal

of Geophysical Research, 98: 16033–16052.

Petronotis, K.E., Gordon, R.G., and Acton, G., 1994. A 57Ma Pacific

plate paleomagnetic pole determined from a skewness analysis of

crossings of marine magnetic anomaly 25r. Geophysical Journal

International, 118: 529–544.

Pitman, W.C., and Heirtzler, J.R., 1966. Magnetic anomalies over the

Pacific-Antarctic ridge. Science, 154: 1164.

Pouliquen, G., Gallet, Y., Patriat, P., Dyment, J., and Tamura, C., 2001.

A geomagnetic record over the last 3.5 million years from deep-

tow magnetic anomaly profiles across the Central Indian Ridge.

Journal of Geophysical Research, 106: 10941–10960.

Raff, A.D., and Mason, R.G., 1961. A magnetic survey off the west

coast of North America, 40-N to 52-N. Bulletin of Geological

Society of America, 72: 1267–1270.

Schouten, H., Klitgord, K., and Gallo, D., 1993. Edge-driven

microplate kinematics. Journal of Geophysical Research, 98:

6689–6702.

Schouten, H., Tivey, M.A., Fornari, D.J., and Cochran, J.R., 1999. The

central anomaly magnetic high: constraints on volcanic construc-

tion of seismic layer 2A at a fast-spreading midocean ridge, the

East Pacific Rise at 9



30–50

N. Earth and Planetary Science Let-

ters, 169:37–50.

Scotese, C.R., 1997. Continental Drift, 7th Edition. Arlington, TX:

PALEOMAP Project, p. 79.

Talwani, M., Windisch, C.C., and Langseth, M.G., 1971. Reykjanes

ridge crest: a detailed geophysical study. Journal of Geophysical

Research, 76: 473–517.

Tivey, M., 1996. Vertical magnetic structure of ocean crust determined

from near-bottom magnetic field measurements. Journal of Geo-

physical Research, 101: 20275–20296.

Tivey, M.A., 2004. Paving the seafloor—brick by brick. Oceanus, 42:

44–47.

Tivey, M.A., Bradley, A., Yoerger, D., Catanach, R., Duester, A., Lib-

eratore, S., and Singh, H., 1997. Autonomous underwater vehicle

maps seafloor. EOS, 78: 229–230.

Tivey, M., Johnson, H.P., Fleutelot, C., Hussenoeder, S., Lawrence, R.,

Waters, C., and Wooding, B., 1998. Direct measurement of mag-

netic reversal polarity boundaries in a cross-section of oceanic

crust. Geophysical Research Letters, 25: 3631–3634.

Vine, F.J., 1966. Spreading of the ocean floor: new evidence. Science,

154: 1405–1425.

Vine, F.J., and Matthews, D.H., 1963. Magnetic anomalies over ocea-

nic ridges. Nature, 199: 947–949.

Vine, F.J., and Wilson, J.T., 1965. Magnetic anomalies over a young

oceanic ridge off Vancouver Island. Science, 150 : 485–489.

Vogt, P.R., Anderson, C.N., and Bracey, D.R., 1971. Mesozoic mag-

netic anomalies, sea floor spreading, and geomagnetic reversals

in the southwestern North Atlantic. Journal of Geophysical

Research, 76: 4796–4823.

Cross-references

Magnetic Anomalies, Marine

Magnetic Surveys, Marine

VOLCANO-ELECTROMAGNETIC EFFECTS

Volcano-electromagnetic effects—electromagnetic (EM) signals gener-

ated by volcanic activity—derive from a variety of physical processes.

These include piezomagnetic effects, electrokinetic effects, fluid vaporiza-

tion, thermal demagnetization/remagnetization, resistivity changes, ther-

mochemical effects, magnetohydrodynamic effects, and blast-excited

traveling ionospheric disturbances (TIDs). Identification of different

physical processes and their interdependence is often possible with multi-

parameter monitoring, now common on volcanoes, since many of these

processes occur with different timescales and some are simultaneously

identified in other geophysical data (deformation, seismic, gas, iono-

spheric disturbances, etc.). EM monitoring plays an important part in

understanding these processes.

Brief history

The identification of electromagnetic field disturbances associated

with volcanic activity has a relatively young history. Initial measure-

ments in the 1950s on Mihara volcano in Japan focused on monitoring

the inclination and declination of the magnetic field while on Kilauea

in Hawaii static electric fields (self-potential, SP) were shown to

reflect the intrusion activity. More robust magnetic and electric field

observations since the 1960s show changes of several tens of nanotesla

and more than a hundred mV km

1

accompany eruptions from volca-

noes (Johnston, 1997; Zlotnicki, 1995). Following eruptions, much lar-

ger changes result from remagnetization processes.

984 VOLCANO-ELECTROMAGNETIC EFFECTS

Physical mechanisms involved

Short-term magnetic and electric field changes during volcanic erup-

tions result primarily from stress changes in volcanic rocks and injection

of hydrothermal/magmatic fluids and gasses. Of particular impor-

tance for research in eruption prediction are changes generated in

the early stages of eruptions by deformation and/or fluid/gas flow. Detec-

tion of these fields provides an independent window into the eruption

process.

The dominant physical processes in play during the final stages

before eruptions are stress-induced modification of magnetization

(piezomagnetism), electric currents produced by fluid flow (electroki-

netic effects), and thermal demagnetization and remagnetization of

volcanic rocks. Other EM signals result from charge generation during

fluid vaporization and the creation of ash. Changes in EM response

also result from modification of the electrical conductivity structure.

During and following eruptions, dramatic EM signals can result from

electric charge generation from rock disintegration, lightning within

the eruption ash cloud, magnetohydrodynamic effects, thermal remag-

netization, stress driven magnetization changes, and coupled TIDs.

It has become increasingly clear, because of the complexity and

interdependence of volcanic processes, that simultaneous multipara-

meter monitoring (e.g., strain, fluid pressure, tilt, ground displacement,

electric fields, magnetic fields, gas emissions, etc.) on volcanoes is

necessary to identify and separate contributions from the different phy-

sical processes. Interpretation of electric and magnetic fields can thus

be best done within the constraints placed by other geophysical data

and vise versa.

Piezomagnetic effects

The magnetic properties of rocks have been shown under laboratory

conditions to depend on the state of applied stress, and theoretical

models for this phenomenon have been developed in terms of single

domain, pseudosingle domain rotation and multidomain wall transla-

tion. The stress sensitivity of magnetization K, defined as the change

in magnetization per unit magnetization per unit stress, can be

expressed in the form

I  Ks  I; (Eq. 1)

where I is the change in magnetization in a body with net magneti-

zation I due to a deviatoric stress s, K the stress sensitivity, typically

has values of about 3  10

3

MPa

1

. The stress sensitivity of induced

and remanent magnetization from theoretical and experimental studies

has been combined with stress estimates from dislocation models of

fault rupture and elastic pressure loading in active volcanoes to calcu-

late magnetic field changes expected to accompany volcanic activity.

These volcanomagnetic models show that moderate scale magnetic

anomalies of several nanoteslas (nT) should accompany moderate to

large volcanic eruptions for rock magnetizations and stress sensitivities

of 1Amperemeter

1

(A m

1

) and 10

3

MPa

1

, respectively (Johnston,

1997).

Electrokinetic effects

Electrokinetic electric and magnetic fields result from fluid flow

through layered volcanic rocks as ions anchored to solid material

induce equivalent ionic charge of opposite sign in the fluids. The fluid

flow itself is thermally and gravitationally driven. Fluid flow in this

system transports the ions in the fluid in the direction of flow, produ-

cing electric potential of several tens of mV km

1

and magnetic fields

of a few nanoteslas.

The current density j and fluid flow rate v in porous media are

found from the coupled equations

j ¼srE 

xz rP



; (Eq. 2)

v ¼

fxzrE



krP



; (Eq. 3)

where E is streaming potential, s is the electrical conductivity of the

fluid, x is the dielectric constant of water,  is fluid viscosity, z is

the zeta potential, f is the porosity, k is the hydraulic permeability,

and P is pore pressure. Unfortunately, many of these parameters are

poorly known within volcanoes.

Electrokinetic models combined with estimates of electrical conduc-

tivity structure within volcanoes have been proposed to explain the

curious “W” form of electric potential commonly observed on active

volcanoes (Ishido, 2004).

Electrical resistivity

Changes in electrical resistivity of volcanic rocks result from changes

in strain, temperature, chemistry, and fluid content. While little theore-

tical research has been done on the details of resistivity changes within

volcanoes and methods to separate the different contributions, there is

no question that resistivity changes occur and are observed during

eruptions (e.g., Yukutake et al., 1990; Zlotnicki et al., 2003). Measure-

ment methods include magnetotellurics (MT) and Schlumberger,

Wenner, DC/DC, and dipole/dipole systems.

Figure V9 Total magnetic field differences between stations SHW, on Mount St. Helens, and remote reference station, PTM, at Portland

Oregon, for the several hours preceding and following the 18 May 1980 catastrophic eruption. Time in UTC (from Johnston et al.,

1981).

VOLCANO-ELECTROMAGNETIC EFFECTS 985