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

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the surface of the subducted plate. Given the resistivity of seawater

(0.3 Om), a porosity of around 1%–2% is sufficient to account for

the observed bulk resistivity of this feature over most of its length. It

is also likely that as the plate subducts the increase in temperature with

depth leads to the release of additional fluid from hydrated minerals.

Low resistivity closer to the surface just to the east of the subduction

trench results from the presence of fluids in the accretionary prism

formed as the bulk of the oceanic sediments are scraped from the sur-

face of the subducted plate. In the eastern part of the model the large

low resistivity region at depths of 20–50 km is interpreted to be due

to water rising from the top of the subducted plate as it is released

by further dehydration reactions.

The low resistivity region in the western part of the EMSLAB

model illustrates the influence of partial melting on resistivity struc-

ture. This feature is interpreted to be due to several percent of partial

melt associated with regional upwelling connected to the Juan de Fuca

Ridge which lies only about 100 km further to the west. An MT study

of the Mid-Atlantic Ridge in Iceland (Oskooi et al., 2005) has likewise

detected low resistivity inferred to be associated with partial melting at

a spreading ridge, and 7% of partial melt has also been calculated as

being necessary to explain the low resistivity of the convection plume

associated with the Hawaiian islands hot spot (Constable and Heinson,

2004).

Whereas, subduction zones mark the boundaries between oceanic

and continental plates, continent-continent collision leads to the rapid

uplift of mountain belts. For example, the Southern Alps of New

Zealand, which result from the collision of the Pacific Plate with the

Australian Plate along the Alpine Fault, are estimated to have under-

gone 20 km of uplift in the last 26 million years. Similarly, the

ongoing collision between the Indian subcontinent and the Asian con-

tinent has led to the uplift of the Himalaya and the Tibetan Plateau.

Major orogenic events such as these appear to be typified by the pre-

sence of crustal low resistivity regions resulting from the presence of

Figure E29 (a) A simple model of a two-phase material consisting

of a regular array of cubic cells of resistivity r

surrounded by a

matrix of resistivity r

. (b) The manner in which the bulk resistivity

of the material varies with the volume fraction of the low

resistivity material.

Figure E30 The role played by fluids in influencing the electrical structure associated with tectonic processes illustrated by

the electrical resistivity model of the Juan de Fuca subduction zone at latitude 44



–46



N. Note the vertical scale changes at

5 and 150 km. (After Wannamaker et al., 1989).

246 EM, TECTONIC INTERPRETATIONS

interconnected fluids. Beneath the Southern Alps low resistivity has

been interpreted as being due to the release and upward migration of

fluids as the development of a crustal root (i.e., a thickening of the

crust beneath the region of uplift) leads to dehydration reactions as

cooler material is pushed to greater depth (Ingham, 1996; Wannamaker

et al., 2002). A somewhat more complex situation is interpreted to

exist beneath the Himalaya and the Tibetan Plateau. Here, significantly

enhanced crustal electrical conductivity, detected over a distance of

nearly 600 km beyond the original collision zone (Wei et al., 2001;

Li et al., 2003), illustrates the difficulty that can occur in differentiat-

ing between reduced resistivity due to partial melting and that due to

the presence of aqueous fluids. Numerical modeling has been used

to place limits on the conductance (the product of the conductivity

and thickness) of the region of anomalously low resistivity beneath

the Tibetan Plateau. The calculated minimum conductance of this area

(6000 S) has then be used, with assumptions concerning the likely

resistivity values for partial melt and saline fluids, to estimate the

volume fractions of these, and the required thickness of the low resis-

tive layer, which would be necessary to explain both the MT and seis-

mic reflection data. The preferred model of a thin layer containing

saline fluids lying above a much thicker layer of partial melt, shows

the importance in such situations of synthesis of EM data with data

from other geophysical techniques.

Many models that have been proposed to explain the physics of the

generation of earthquakes have concentrated on the role played by

fluids within fault zones. The formation of a fractured, cataclastic zone

along an active fault allows percolation of fluids into the damaged

zone, and there has been considerable debate over the possibility that

the weakness of many faults may therefore result from the presence

of high pressure fluids in the fault zone. As a consequence there have

been a number of EM studies aimed at detecting and delineating con-

ductive zones associated with faults, and which have interpreted these

in terms of the presence of fluids. For example, a series of MT trans-

ects of the San Andreas Fault (Unsworth et al., 1999, 2000; Bedrosian

et al., 2002) have suggested a correlation between the presence of an

electrically conductive zone along the fault trace and whether the fault

zone is locked or not. The existence of a significant fault zone conduc-

tor (FZC) on a creeping segment of the San Andreas at Hollister has

been associated with a higher fluid content within the fault zone.

Whether the presence of fluid in the fault zone facilitates creep of

the fault, or is itself a result of creep leading to the opening of

fractures, remains somewhat uncertain. However, in contrast, locked

segments of the fault zone exhibit only minor FZCs. At Parkfield,

the most intensively studied segment of the San Andreas Fault because

of its regular recurrence interval for earthquakes, the base of a moder-

ate FZC coincides with the depth beneath which microearthquake

activity is observed. Estimates of the fluid resistivity suggest that to

achieve the measured bulk resistivity within the FZC at Hollister and

Parkfield the porosity must be between 10% and 30%.

Similar conductive zones have been noted to be associated with

other major transform faults, although some prominent faults do not

exhibit FZCs (e.g., Ritter et al., 2003). Nevertheless, the potential

importance of understanding the relationship between fluids, FZCs,

and the genesis of earthquakes is illustrated by MT measurements

(Tank et al., 2005) on the North Anatolian Fault, which marks the

boundary between the Eurasian and Anatolian Plates in northern

Turkey. These show that the hypocenters of both the mainshock and

aftershocks of the M

¼ 7.4, 1999 Izmit earthquake were located

close to a boundary between resistive and conductive zones.

One other obvious manifestation of tectonism is represented

by volcanic activity and there have been numerous EM studies of

active volcanoes (e.g., Ogawa et al., 1998; Muller and Haak, 2004).

Although, the low resistivity of molten rock (0.1 Om) suggests that

the magma chamber of a volcano ought to present a highly significant

electrical target for EM measurements, there are in reality a number of

complicating factors which make detection of magma chambers diffi-

cult. These include the existence of zones of low resistivity which

are the result of either thermal alteration of rocks or the presence

of hydrothermal systems associated with volcanic vents. These low

resistivity regions normally occur close to the surface and therefore,

because EM fields attenuate with depth more rapidly in a low resistiv-

ity material (the skin-depth effect), have the effect of decreasing the

ability to resolve deeper structural features. Other complicating factors

arise from the three-dimensional nature of the electrical structure

associated with volcanoes, and the fact that measurements may also

be significantly affected by topography.

Malcolm Ingham

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ric imaging of the creeping segment of the San Andreas Fault

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Constable, S., and Heinson, G., 2004. Hawaiian hot spot swell struc-

ture from seafloor MT sounding. Tectonophysics, 389: 111–124.

Ingham, M., 1996. Magnetotelluric soundings across the South Island

of New Zealand: electrical structure associated with the orogen

of the Southern Alps. Geophysical Journal International, 124:

134–148.

Li, S., Unsworth, M.J., Booker, J.R., Wei, W., Tan, H., and Jones,

A.G., 2003. Partial melt or aqueous fluid in the mid-crust of south-

ern Tibet? Constraints from INDEPTH magnetotelluric data. Geo-

physical Journal International, 153: 289–304.

Minarik, W.G., and Watson, E.B., 1995. Interconnectivity of carbonate

melt at low melt fraction. Earth and Planetary Science Letters,

133: 423–437.

Muller, A., and Haak, V., 2004. 3-D modelling of the deep electrical

conductivity of Merapi volcano (Central Java): integrating magne-

totellurics, induction vectors, and the effects of steep topography.

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Ogawa, Y.N., Matsushima, H., Oshima, H., Takakura, S., Utsugi, M.,

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Oskooi, B., Pedersen, L.B., Smirnov, M., Arnason, K., Eysteinsson,

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Mid-Atlantic Ridge deduced from MT data in SW Iceland. Physics

of the Earth and Planetary Interiors, 150: 183–195.

Ritter,O.,Ryberg,T.,Weckmann,U.,Hoffmann-Rothe,A.,Abueladas,A.,

Garfunkel, Z., and DESERT Research Group, 2003. Geophysical

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netotelluric imaging of the fault rupture area of the 1999 Izmit

(Turkey) earthquake. Physics of the Earth and Planetary Interiors,

150: 213–225.

Unsworth, M., Egbert, G., and Booker, J., 1999. High-resolution elec-

tromagnetic imaging of the San Andreas fault in Central California.

Journal of Geophysical Research, 104 : 1131–1150.

Unsworth, M., Bedrosian, P., Eisel, M., Egbert, G., and Siripunvaraporn, W.,

2000. Along strike variation s in the elect rical structure of the San

Andreas Fault at Parkfield, California. Geophysical Research Letters

27: 3021–3024.

Wannamaker, P.E., Booker, J.R., Jones, A.G., Chave, A.D., Filloux,

J.H., Waff, H.S., and Law, L.K., 1989. Resistivity cross section

through the Juan de Fuca subduction system and its tectonic impli-

cations. Journal of Geophysical Research, 94: 14127–14144.

Wannamaker, P.E., Jiracek, G.R., Stodt, J.A., Caldwell, T.G., Gonza-

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tion and pathways beneath an active compressional orogen, the

EM, TECTONIC INTERPRETATIONS 247

New Zealand Southern Alps, inferred from magnetotelluric

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Wei, W., Unsworth, M., Jones, A.G., Booker, J., Tan, H., Nelson, D.,

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Cross-references

EM, Regional Studies

Magnetotellurics

ENVIRONMENTAL MAGNETISM

Introduction

Environmental magnetism, which has rapidly developed into an estab-

lished science from its beginnings some 25 years ago, involves the

application of magnetic methods to environmental materials, in order

to measure their magnetic properties. The environmental materials in

question encompass soils, sediments, dusts, rocks, organic tissues,

and man-made materials. Their magnetic properties vary according

to the mineralogy, concentration, grain size, morphology, and compo-

sition of the magnetic minerals present. These (nondirectional) mag-

netic properties can act as sensitive recorders of environmental and

climatic information, both for the present day and through geological

time. Paleo- and rock magnetists use magnetic measurements to iden-

tify which magnetic minerals were responsible for acquiring a record

of the Earth's magnetic polarity at their time of formation. Environ-

mental magnetists seek to identify the causal links between magnetic

properties and ambient climatic, environmental and postdepositional

processes. The magnetic properties can then be used, inversely, to

identify natural and human-induced changes in climate and environ-

ment on both site-specific and regional scales and over current and

geological timescales.

Environmental magnetism has become a valued and widely applied

methodology because: magnetically distinctive assemblages of mag-

netic grains—dominantly iron oxides, oxyhydroxides, and sulfides—

occur virtually ubiquitously throughout the natural and built

environment; these assemblages vary according to their source and

depositional history; and sensitive magnetic measurements (i.e., to con-

centrations of less than 1 part per million for weakly magnetic minerals,

like hematite, and less than 1 part per billion for strongly magnetic

minerals, like magnetite) can be made relatively cheaply and rapidly

compared with other types of mineralogical analysis.

Historical material

As early as 1926, magnetism was used to examine environmental pro-

cesses, in a glacial setting. Gustav Ising measured the magnetic sus-

ceptibility and natural magnetic remanence of annually laminated lake

sediments from Sweden. He found that the magnetic properties of the

sediments varied seasonally and with distance from their ice margin

source. The lake sediment layers deposited in the springtime were much

more magnetic than those deposited in winter. Ising interpreted these

magnetic changes as reflecting the high specific weight of magnetite

grains and thus their response to different rates of flow of glacial melt-

water. More magnetite grains were carried at higher rates of (springtime)

flow. More than 40 years later, John Mackereth measured magnetic

properties of sediments from Windermere, in the English Lake District.

His work sparked new interest in the sources of magnetic minerals

found in lake sediments and their surrounding catchment soils.

Since the mid-1970s onward, magnetic analyses have been applied

to the sediment records of over 100 lakes worldwide. The last 20 years

have seen the application of environmental magnetism to solve diverse

and global-scale research questions, including: high-resolution correla-

tion of sedimentary units (in marine, lacustrine, and marine-to-terrestrial

settings); identification of the flux and timing of detrital inputs

(ice-rafted debris, aeolian dust) in the deep-sea sedimentary record;

unraveling of postdepositional diagenetic processes in sediments; iden-

tification of ferrimagnet formation, and destruction in different types

of soils; discovery of sedimentary magnetic mineral contributions by

magnetotactic bacteria; identification of natural magnets in a range

of different organisms (insects, fish, humans); use of magnetism as a

quantitative proxy for paleorainfall in the famous loess/soil sequences

of East Europe and East Asia; archaeological prospection; sourcing

of fluvial sediments; and mapping of particulate pollution sources

and loadings, especially from vehicles (Figure E31). More than 130

magnetic laboratories worldwide are presently engaged in environmental

magnetism research.

Essential concepts and applications

The environmental magnetic approach can briefly be summarized as

the investigation of magnetic mineralogies in natural and anthropo-

genic samples by measurement of both their induced and remanent

responses to a series of artificially induced magnetic fields. The speed

of magnetic measurements, compared with other types of mineralogi-

cal analysis, and their nondestructive nature, enables measurement

of large numbers of samples, a key advantage in obtaining high-

resolution (spatial and/or temporal) environmental data. For any envir-

onmental sample, a large number of different magnetic parameters

could be measured in order to characterize its magnetic properties and

investigate its magnetic mineralogy. In practice, a reasonably small

number of magnetic measurements can often provide sufficient infor-

mation to identify its major magnetic constituents (Figure E32). Table

E1 and Figure E33 (Maher et al., 1999) summarise some of the mag-

netic parameters used routinely in environmental magnetism and typi-

cal values for a range of magnetic minerals. Figure E34 shows how

magnetic ratio parameters can be used to discriminate between differ-

ent grain sizes in ferrimagnets. Environmental materials normally con-

tain mixtures of magnetic minerals and magnetic grain sizes. Work is

in progress to extend the capability of magnetic methods in order to

identify, quantify, and ascribe specific environmental histories/origins

to the individual magnetic components within such mixed samples.

In the meantime, independent and complementary analyses are also

often used, in order to more fully characterize sample mineral proper-

ties and compare the magnetic data with other environmental proxies

(such as oxygen isotope ratios or sample geochemistry).

The carriers of the magnetic properties of rocks are normally the iron

oxides, magnetite, titanomagnetite, maghemite, and hematite. Until

recently, because of their significance as paleomagnetic recorders, the

titanium-substituted magnetites and maghemites, formed at high tem-

peratures in igneous rocks, were a particular focus of intensive rock mag-

netic research. Additional rock magnetic attention was also given to

hematite, formed at lower temperatures in terrestrial settings by oxida-

tion of iron minerals, during weathering and transport, or postdeposi-

tionally by oxidizing pore fluids. Weathering, erosion, and transport of

these lithogenic minerals can provide one pathway for the supply of

magnetic minerals to a range of environmental materials (soils, sedi-

ments, dusts). However, a key advance brought about in environmental

magnetism has been the identification of a diverse range of processes

and pathways of ferrimagnet formation at or near the Earth's surface,

as described below.

Direct bacterial formation of magnetite and greigite, of submicrometer

grain size, has been discovered (e.g., Blakemore, 1975; Bazylinski,

1991). A range of microaerobic and anaerobic magnetogenic species

actively assimilate iron from the extracellular environment, to form

248 ENVIRONMENTAL MAGNETISM

distinctive, intracellular chains of ferrimagnetic crystals, some exhibiting

unique crystal morphologies (e.g., bullet, boot, and blade shapes). These

biogenic crystals normally fall within a narrow grain size range, from

20 to 50 nm, and behave as interacting, stable single-domain grains,

resulting in strong and stable magnetic behavior. As a result, the torque

of the Earth's magnetic field is transferred effectively to the bacterial

cell, so that when the bacteria swim, they move along the Earth's field

lines. Magnetogenic bacteria, and/or their bacterial magnetofossils (Fig-

ure E35), have been found in diverse sedimentary environments, includ-

ing marine (Petersen et al., 1986), lake (Spring et al., 1993), brackish

(Stolz, 1992), and saltmarsh (Sparks et al., 1989) sediments, ocean

waters (Bazylinski et al., 1988), soils (Fassbinder et al., 1990), rivers

(Bazylinski, 1991), and sewage ponds (Moench and Konetzka, 1978).

In terms of population numbers, Petersen et al. (1986) estimated for

some lake sediments in South Germany a population density of 10

mag-

netogenic bacteria ml

1

. In pure culture, the amount of magnetite pro-

duced by magnetogenic bacteria has been estimated at about 0.2 g per

10 g (wet weight) of bacteria (equivalent to 10

bacterial cells; Fran-

kel, 1987). Other organisms have also been found to synthesize

ferrimagnets, including insects (bees, termites, butterflies), fish, birds,

and humans.

Nanocrystalline magnetite/maghemite has also been found to form

inorganically, in situ, and at ambient temperatures (<50



C) within

well-drained (generally oxidizing), buffered soils (pH range 5–8),

as a result of release of iron by weathering and subsequent repeated

redox cycles (Mullins, 1977; Maher and Taylor, 1988). Magnetite

is a mixed Fe

2þ

/Fe

3þ

iron oxide; its in situ formation (in the trace

amounts responsible for the observed magnetic enhancement in soils)

requires the initial presence in the soil matrix of some Fe

2þ

. This path-

way of ferrimagnet formation may be mediated by the action of iron-

reducing bacteria, which use iron—in the absence of oxygen—as the

terminal electron acceptor in their digestion of organic matter. They

release the resultant Fe

2þ

into the extracellular environment, where it

may partially oxidize to form magnetite (Lovley et al., 1987) and/or

by surface adsorption cause other soil iron oxides and oxyhydroxides

to dissolve, with possible magnetite formation via a green rust phase

(Tamaura et al., 1983). (Action of such Fe-reducing bacteria may also

contribute to authigenic formation of magnetite in suboxic sediments

of the deep-sea floor, see below.) Unlike bacterial magnetite, which

is biologically constrained in crystal size and shape by the intracellular

membrane structures into which it is precipitated, extracellular, pedo-

genically formed magnetite is variable in grain size (1– 500 nm) and

does not display the unique crystal morphologies observed in bacterial

magnetite chains. DNA screening of magnetically enriched topsoils has

shown that populations of magnetotactic bacteria are too low (<10

bacteria g

1

) to account for observed soil ferrimagnetic concentrations

(Dearing et al., 2001), whereas the action of one iron-reducing bacter-

ium can mediate the extracellular formation of hundreds of ferrimag-

netic grains.

So-called “botanical,” nanocrystalline magnetite (1–50 nm) has

recently been identified (by transmission electron microscopy and

electron diffraction analyses) in grass plant cells. The magnetite is

located within the inorganic cores of the plant iron storage compound,

phytoferritin (Gajdardziska-Josifovska et al., 2001). Upon breakdown

of the organic material, this botanical pathway may constitute another

(as yet unquantified) contributor to soil ferrimagnets.

Depending on sedimentation rate and organic matter content, deep-

sea sediments can contain authigenic and diagenetic magnetic oxides

and/or sulfides, in addition to any detrital magnetic minerals. Gener-

ally, as the amount of sedimentary organic matter increases, micro-

organisms use available oxidants (electron acceptors) in order to

metabolize the organic material, with iron oxides being reduced once

available nitrate and manganese oxides have been used up. Finally,

when the iron oxides too have been reduced, sulfate is used as the oxi-

dant, with the potential to form some magnetic iron sulfides (such as

greigite), if this process does not proceed through to formation of

pyrite (e.g., Snowball and Thompson, 1988; Tric et al., 1991). Thus,

magnetic iron oxides and sulfides can form and dissolve in deep-sea

sediments, depending on sedimentary geochemical conditions (Karlin

et al., 1987; Langereis and Dekkers, 1999). In turn, the latter are often

influenced by changes in climate, especially in generating variations in

amounts of organic matter and rate and type of sediment delivery.

In addition to these low-temperature sources of ferrimagnets, two

high-temperature pathways have also been identified. Combustion

Figure E31 Magnetic minerals in the environment and examples of major areas of application of environmental magnetism

(adapted from Thompson and Oldfield, 1986).

ENVIRONMENTAL MAGNETISM 249

Figure E32 A flowchart indicating how a relatively small number of magnetic measurements can be used to discriminate between the

major magnetic minerals found in environmental samples (reproduced with permission from Maher et al., 1999).

Table E1 Typical magnetic properties of natural minerals (reproduced with permission from Maher et al., 1999)

w ARM SIRM

Mineral (



C) (Am

1

)(mm

1

) (mAm

1

)(Am

1

)

Magnetite (soft) 575 92 560 18 9

Magnetite (hard) 575 92 400 110 22

Titanomagnetite (soft) 200 24 170 80 7

Titanomagnetite (hard) 200 24 200 480 12

Hematite 675 0.5 0.6 0.002 0.24

Ilmenohematite 100 30 25 480 8

Greigite 300 20 120 110 11

Pyrrhotite 300 17 50 80 4.5

Goethite 150 0.5 0.7 0.005 0.05

Iron 770 220 2000 800 80

Paramagnet 1/T – 10 0

Diamagnet const – 0.006 0 0

Note: T

, Curie temperature or thermomagnetic behaviour; M

, saturation magnetization; w, initial AC susce ptibility; ARM, anhysteretic remanence grown in a 100 mT peak

alternating field biased with a 0.1 mT DC field; SIRM, isothermal remanence grown in a 1 T DC field.

250 ENVIRONMENTAL MAGNETISM

of organic matter in soils can form ferrimagnets from other soil iron

oxides and oxyhydroxides (Longworth et al., 1979), while fossil fuel

combustion also releases magnetite-bearing, pollutant particles into

the atmosphere (Puffer et al., 1980; Hunt et al., 1984, and see below).

The magnetic properties of environmental samples are largely domi-

nated by the strongly magnetic ferrimagnets. However, high-field mag-

netic experiments (i.e., using DC fields of >2 T) can be used to

selectively investigate the weakly magnetic minerals goethite and

hematite, even when ferrimagnet concentrations are high. In samples

deficient both in ferrimagnets and imperfect antiferromagnets, the

weak paramagnetism of iron-bearing silicates and the negative, dia-

magnetic susceptibility of carbonate can be significant (Hounslow

and Maher, 1999).

Applications

In modern process studies, the magnetic properties of contemporary

soils have been used to identify soil-forming processes (Maher,

1986, 1998), to relatively date weathering surfaces (Pope, 2000), to

“fingerprint” different soil types and soil horizons in order to monitor

soil erosion and identify sediment sources (e.g., Walling et al., 1979;

Dearing et al., 1985; Oldfield et al., 1985), and to identify relation-

ships between modern climate and modern soil magnetic properties

(Maher and Thompson, 1995; Han et al., 1996), in order to make

paleoclimatic estimates from paleosols (“fossil” or buried soils). This

latter approach has been applied to the famous loess/paleosol

sequences of the Chinese Loess Plateau, which span quasicontinuously

the entire Quaternary period. These alternating windblown loess and

soil layers represent colder (and drier) climate stages, when loess

deposition proceeded with little soil formation, and warmer, wetter

stages, when loess was still deposited but weathering and soil forma-

tion also occurred to varying degrees. The magnetic susceptibility

values of the soil layers are 3–5 times higher than the intervening

loess units, reflecting in situ pedogenic formation of ferrimagnets (Fig-

ure E36). To isolate the in situ, soil-formed magnetic signal, the pedo-

genic susceptibility (w

ped

) was defined for each soil as the maximum

susceptibility value (w) of the B horizon minus the w of the parent

loess. Modern soils, across the Chinese Loess Plateau and the loessic

Russian steppe display strong positive correlation between pedogenic

w (and other magnetic parameters, such as frequency-dependent sus-

ceptibility, anhysteretic remanence), and annual rainfall (Heller et al.,

1993; Maher et al., 1994, 2002). Applying this modern rainfall/

magnetism couple to the paleosusceptibility values of the buried soils

Figure E33 A magnetic hysteresis (magnetization vs field) loop for

a typical natural sample containing a mixture of ferrimagnetic and

paramagnetic minerals and some of the magnetic parameters

routinely used in environmental magnetism (reproduced with

permission from Maher et al., 1999).

Figure E34 The grain size dependence of selected magnetic properties for magnetite: magnetic remanence ratio (the susceptibility of

the anhysteretic remanence/the saturation remanence) and demagnetization (the median destructive field of the anhysteretic

remanence) data for discrimination of magnetic grain size in ferrimagnets (from Maher, 1988).

ENVIRONMENTAL MAGNETISM 251

enables calculation of paleorainfall values through the Quaternary.

Higher average rainfall is identified for the early to mid-Holocene than

at the present day, indicating stronger northward and westward incur-

sion of the rain-bearing summer monsoon winds across the SE Asian

continent. Conversely, during glacial stages, rainfall decreased by up

to 50%. Kukla et al. (1988) contest this magnetism/rainfall relation-

ship, and suggest that the loess/paleosol magnetic variations are con-

trolled by variations in the flux of the weakly magnetic loess—loess

flux is higher, and soil magnetic properties weaker, in the northwest

of the Loess Plateau. But the loessic Russian steppe, where there is

no loess deposition at the present day, also demonstrates the rainfall/

magnetism correlation observed in China, indicating the key influence

of climate, rather than dust flux, in controlling soil magnetic properties

in these areas. Finally, the loess/paleosol magnetic variations correlate

with the deep-sea oxygen isotope record, showing that the Asian mon-

soon systems are interlinked with the processes of midlatitude ice

sheet growth and decay.

Increasingly, magnetic measurements are also being applied to pre-

Quaternary paleosols, in order to identify their soil-forming processes

and environments. For example, the magnetic properties of paleosols

of Permian age appear to record glacial-interglacial fluctuations in

low-latitude Pangea (e.g., Tramp et al., 2004).

Soils and weathering bedrock can also be considered as sediment

precursors. If they display distinctive magnetic properties, preserved

through both erosion and subsequent deposition as sediment, then it

may be possible to construct soil-sediment magnetic linkages and

quantitatively identify changes in rates of flux of those sources through

time. For instance, Caitcheon (1993) demonstrated that magnetism can

be used to identify sources of suspended and bedload sediments at

stream tributary junctions. The tributary contributions to the resultant

downstream mix were also quantifiable. Further developments in

quantitative sediment source ascription encompass multivariate statisti-

cal unmixing of sediment magnetic properties against potential source

magnetic properties, using various types of algorithms, including mul-

tiple regression, maximum likelihood unmixing (Shankar et al., 1994)

and linear programming (Walden et al., 1997). For example, Shankar

et al. (1994) estimated that natural catchment sources contributed

>97% of stream bedload upstream of an iron ore mine, but that mine

waste contributed 50% of the downstream bedload. An UK-based

study of suspended sediment sources (Gruszowski et al., 2003), based

on magnetic, geochemical, and radionuclide signatures (in the catch-

ment of the River Leadon), was able to estimate relative contributions

of 43% from subsurface (subsoil and channel bank) sources, and

27% from topsoil (arable and grassland) sources. The most recent

development in tracing of fluvial sediment sources is based on the

magnetic signatures of magnetic minerals carried as inclusions within

host silicate grains (Hounslow and Morton, 2004). By removing any

discrete magnetic grains (through premeasurement acid treatment),

Figure E35 Transmission electron micrograph of bacterial

magnetofossils, some with the unique “bullet” and “boot”

morphologies, extracted from Cretaceous chalk, Culver Cliff,

United Kingdom (photo: M. Hounslow, reproduced with

permission from Maher et al., 1999).

Figure E36 The interbedded loess (pale color) and paleosols (darker layers) of the Chinese Loess Plateau, at Luochuan (central Loess

Plateau) together with magnetic susceptibility (in units of 10

–8

1

) and sedimentary logs (S

, modern soil; S

, Paleosol 1, L

Loess 1) for the arid western plateau, and the increasingly humid central and southeastern plateau. Note that the loess/soil sequences

are thicker in the west and have increasingly high magnetic susceptibility values toward the southeast.

252 ENVIRONMENTAL MAGNETISM

the sample magnetic properties reflect only magnetic inclusion signa-

tures. As the magnetic inclusions occur within host particles, they

are not subject to any hydraulic sorting effects and are protected from

any postdepositional, diagenetic alteration/erasure.

Providing a temporal dimension on historical and/or geological

timescales, lake, estuarine, and deep-sea sediments can form natural

records both of detrital fluxes and productivity in their aquatic commu-

nities. The measured magnetic properties of such sediments can thus

reflect changes in allogenic inputs through time and/or authigenesis

and/or postdepositional diagenesis. Magnetic methods are thus often

embedded within other, multiproxy analyses in order to discern the

relative significance of these possible inputs. Quantitative “unmixing”

of sediment magnetic properties in terms of possible sediment sources,

as described above, is increasingly applied in lake sediment studies

(e.g., Thompson, 1986; Yu and Oldfield, 1993; Egli, 2004). In the

context of allogenic influx, and at the land catchment-scale, much

environmental magnetic work has focused on using the magnetic prop-

erties of lake sediments to construct source-sediment linkages (as

above) but also to identify sedimentation rates through time (e.g.,

Bloemendal et al., 1979; Dearing and Foster, 1986; Loizeau et al.,

1997), often related causally to catchment disturbance, whether natural

or anthropogenic. Lake sediments can receive detrital inputs from flu-

vial supply and from aeolian flux, of dust, volcanic ash, and pollutants.

Magnetic oxides can form in situ, through the direct bacterial actions

of magnetogenic bacteria and/or the extracellular mediation of the

iron-reducers (as above). In organic-rich conditions, dissolution of

magnetic minerals may occur (e.g., Hilton and Lishman, 1985) while,

with a source of sulfate, magnetic iron sulfides may be precipitated

(e.g., Snowball and Thompson, 1988; Roberts et al., 1996). Hence,

the magnetic properties of lake sediments need to be evaluated on

a site-specific basis, and with appropriate age control, in order to

identify the key magnetic sources through time, and, inversely, to thus

reconstruct past environmental and/or climatic changes. Dearing

(1999) provides a review of lake sediment magnetic records for the

Holocene period. Where fluvial detrital flux of magnetic minerals

can be shown to be of primary importance, changes in hydrological

regime, including rainfall (Foster et al., 2003), stream power (e.g.,

Dearing et al., 1990), snow melt (Snowball et al., 2002), flood fre-

quency and magnitude (Thorndycraft et al., 1998) may be recorded

in lake sediment sequences. Conversely, changes in sediment source

(e.g., Reynolds et al., 2004) or changes in source magnetic properties,

such as results from fire-induced magnetic mineral formation (e.g.,

Rummery, 1983), may be identifiable. In historical times, anthro-

pogenic pollution particles have been a dominant contributor to the

magnetic properties of some lake and estuarine sediments (e.g., Locke

and Bertine, 1986; Battarbee et al., 1988; Kodama et al., 1997).

As for lake sediments, the magnetic properties of deep-sea sedi-

ments can vary as a result of changes in detrital flux (e.g., aeolian dust,

iceberg-rafted debris, and bottom-water transport), bacterially pro-

duced or mediated mineral authigenesis of magnetic oxides or sulfides,

and/or postdepositional diagenesis. Hence, environmental magnetic

properties of marine sediments can also act as proxy indicators of

environmental and climatic change, over geological timescales. Sedi-

ment redox cycles, influenced by hydrological runoff, sedimentation

rates, and organic matter content, can be identified from sediment

magnetic properties, as magnetic minerals form, dissolve, and re-form

through the diagenetic zone. Sediment magnetic properties in restricted

basins like the Mediterranean are very sensitive to even small changes

in climate, effectively amplifying the variations in such regions (e.g.,

Langereis and Dekkers, 1999). Conversely, in the oxic/suboxic sedi-

ments of the Atlantic, sediment magnetic properties dominantly reflect

differences in detrital sediment sources (although where detrital flux is

low, bacterial magnetofossils may constitute a significant contributor

to sediment magnetism). Schmidt et al. (1999), for instance, identified

major sediment sources at the present-day for the South Atlantic,

including the riverine runoff from the Amazon Basin and dust flux

from the African continent. For the North Atlantic, statistical analysis

of a multivariate magnetic data set representing >300 surface sedi-

ments across the ocean identified the present spatial distribution of

iceberg-rafted debris, dominantly from Greenland and the Labrador

Sea/Baffin Bay regions, and two distinctive North African dust plumes

(Watkins and Maher, 2003). Mapping and sourcing of iceberg-rafted

debris for past glacial stages enables reconstruction of ocean surface

currents and testing of ocean circulation models. With regard to

aeolian fluxes, robust identification of dust flux to the deep-ocean sedi-

mentary records during past glacial and interglacial stages is important,

as dust may play an active role in climate change. Dust can change the

radiative properties of the Earth's atmosphere but can also “fertilize”

iron-deficient areas of the world's oceans, such as the Southern Ocean,

stimulating marine productivity and resulting in drawdown of carbon,

with resultant reduction in atmospheric CO

levels (Martin, 1990).

Dust from the arid/semiarid zone of Africa is characterized by the pre-

sence of hematite and/or goethite, weakly magnetic iron oxides with

distinctive high-field magnetic properties (Robinson, 1986; Watkins

and Maher, 2003). It has thus been possible to identify the flux and

provenance of dust by magnetic measurements of suitable deep-sea

and lake sediments (e.g., Maher and Dennis, 2001; Kissel et al.,

2003; Larrasoaña et al., 2004). Other applications of magnetic mea-

surements to deep-sea sediments include: core correlation; tracing of

turbidite flow paths (Abdeldayem et al., 2004), and identification

of hydrothermal alteration zones (Urbat and Brandau, 2003).

Current investigations, controversies, and gaps in

current kn owledge

Recent innovations in environmental magnetism include the use of

magnetic measurements of roadside tree leaves and soils as natural

pollution monitors (e.g., Matzka and Maher, 1999; Knab et al.,

2000; Hanesch et al., 2003); magnetic mapping and in situ measure-

ment of magnetic susceptibility in order to identify levels of contami-

nants in lake sediments (Boyce et al., 2004); magnetic identification of

dust flux to ice cores (Lanci et al., 2001); and identification and corre-

lation of volcanic tephras (e.g., Andrews et al., 2002). Development of

new measurement techniques continues, such as the use of first-order

reversal curves (FORC) diagrams to identify magnetic mineral compo-

sitions and grain sizes (e.g., Pike et al., 2001; Carvallo et al., 2004).

FORC may be suitable for addressing the problem of investigating

the weakly magnetic minerals, hematite and goethite even when mag-

netite is the dominant mineral (Muxworthy et al., 2004). Remanence

measurements at high fields (>2 T) and low temperature (77 K) may

also provide a new means of defining the concentration and grain size

of these minerals in this context (Maher et al., 2004). Finally, to

achieve standardized accuracy as well as precision, interlaboratory

calibration is an ongoing research activity internationally (e.g.,

Sagnotti et al., 2003).

Summary/conclusions

The ubiquity and environmental sensitivity of the magnetic mineralo-

gies present in soils, sediments, and dusts can provide a record

of past and present environmental and climatic change and process.

The speed and sensitivity of magnetic measurements, and their demon-

strated range of environmental applications, has led to their application

to a range of key environmental and climatic questions and problems,

ranging from site-specific to global in scale. Environmental magnetism

has thus become an established discipline, which continues to gain

in strength from improved measurement capabilities, quantification,

and interaction with other, complementary environmental analyses

and proxies.

Barbara A. Maher

ENVIRONMENTAL MAGNETISM 253

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