Variations in the Intensity of the Earth's Magnetic Field

VARIATIONS IN THE INTENSITY OF THE EARTH’S MAGNETICFIELD

J. A. JACOBSInstitute of Geography and Earth Sciences, University of Wales, Aberystwyth SY23 3DB, U.K.

Abstract. The Earth’s main magnetic field can be approximated by an axial, geocentric dipole. Theremaining non-dipole field is much smaller and is a regional rather than a global feature – quitelarge changes can occur in a few ka. This review is concerned with changes in the dipole componentof the geomagnetic field, and one of the problems is in separating the non-dipole from the dipolecontributions to the field. Unlike the many determinations of the direction of the Earth’s magneticfield in the past (which have led to fundamental contributions to our understanding of plate tectonicsand shown that the field can on occasion reverse its polarity), estimates of the intensity of the fieldare comparatively few, especially before the Holocene. This is mainly the result of experimentaldifficulties in obtaining reliable measurements of the field. These problems are discussed in somedetail and are followed by a short account of archaeomagnetic intensities and results from Hawaiiwhere many of the first determinations were obtained. Measurements for∼ the last 100 ka from bothlavas and lacustrine and oceanic sediments are reviewed and results from different areas compared.An asymmetric saw-tooth pattern has been observed in some of the records over the last few Ma,and this rather controversial question is discussed. Finally an account is given of the far more limiteddata on palaeointensities in earlier times.

A short discussion is given of the interpretation of coherent short wavelength variations which areobserved in many marine magnetic profiles. Although short reversals of the field may be responsiblefor some of these “tiny wiggles”, it is more likely that in general they are the result of changes in thestrength of the Earth’s magnetic field.

Keywords: Earth, magnetic field, palaeointensity

1. Introduction

The intensity and direction of the Earth’s magnetic field vary from place to placeacross the Earth and also show temporal changes. These range from variations on atimescale of seconds to secular variations on a timescale of hundreds of years andon a longer timescale to complete reversals of polarity. The short period transientvariations produce no large or enduring changes in the Earth’s magnetic field andarise from causes outside the Earth. They will not be considered here. A sphericalharmonic analysis of the Earth’s magnetic field shows that it is of internal originand that to a first approximation it can be represented by a geocentric axial dipole.Although much weaker, the non-dipole components show more rapid changes –the timescale of the non-dipole changes is measured in decades and that of thedipole in centuries. However the non-dipole field is much more complex than the

Surveys in Geophysics19: 139–187, 1998.© 1998Kluwer Academic Publishers. Printed in the Netherlands.

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dipole field. Bloxham and Jackson (1992) and Hulot and Le Moüel (1994), usinghistorical magnetic data for the last 300 years, have shown that there are largenon-dipole variations on time scales of centuries. Gubbins and Kelly (1993) havemodelled palaeomagnetic directions from the last 2.5 Ma using the same methodsas developed by Bloxham et al. (1989) for modern and historical data. They founda striking resemblance in the radial component of the magnetic field at the core-mantle boundary – in particular two lobes of high field in the northern hemispheresituated over Arctic Canada and Siberia. Similarities in the southern hemisphereare much less clear and the field appears to have been smoothed out. Gubbins andKelly suggest that the current rapid secular variation has persisted in this region forseveral million years to produce a nearly axisymmetric pattern.

Studies of the field over the last few hundred years have shown that the sec-ular variation is a regional rather than a planetary phenomenon and considerablechanges can take place in the general pattern even within twenty years. For the pur-pose of comparing data from sampling sites at different latitudes, one can calculatethe equivalent dipole moment which would have produced the measured intensityat the calculated palaeolatitude of the sample – such a dipole moment is called aVirtual Dipole Moment (VDM) or Virtual Axial Dipole Moment (VADM) if noknowledge of the magnetic inclination is available and the dipole axis is assumedto be the axis of rotation (the geographic axis). It is not possible to identify thatportion of the observed palaeointensity (and therefore of the VDM) which is due tothe non-dipole components. According to Tric et al. (1994) the differences betweenVADMs and VDMs are unlikely to exceed±20 per cent unless the inclination isa very long way from the expected dipolar value. It is the purpose of this paper toreview our knowledge of the variations in the intensity of the Earth’s dipole field,and a major problem is the removal of the effects of the palaeo-secular variation(PSV) – over the past 10 ka, the VDM has varied by more than a factor of two.

2. Palaeointensity Measurements

Palaeointensity measurements of the Earth’s magnetic field are much more diffi-cult to make than palaeo-directions and palaeointensities are very scarce beforethe Holocene. In order to obtain the intensity of the ancient field that magnetizeda sample, the natural processes by which the magnetization was acquired mustbe duplicated in the laboratory. This is possible only for volcanic rocks and firedarchaeological material such as pottery which acquire a thermo-remanent magneti-zation (TRM) as they cool from above the Curie point of their magnetic minerals toambient temperature in the Earth’s magnetic field. For weak magnetic fields like theEarth’s, the natural remanent magnetization (NRM) of a sample is proportional tothe strength of the magnetic field in which it is produced. Thus from a comparisonof the NRM with a laboratory TRM in a known field, the ancient field strengthmay be determined provided that the primary magnetic minerals have not altered

VARIATIONS IN THE INTENSITY OF THE EARTH’S MAGNETIC FIELD 141

Figure 1.NRM-TRM diagram showing an example of the “kinks” that often signal the beginningof the magnetochemical changes during the Thelliers’ experiments. Circles are NRM-TRM pointscalculated for each double heating at the temperature indicated near each point. Lines connect eachp

TRM check (triangles) to the NRM-TRM point corresponding to the maximum temperature to whichthe specimen was heated before performing the check. Solid circles are NRM-TRM points used todetermine the least squares line, shown as the heavy, solid line in the figure, open circles are pointsnot used in the calculations. (after Mankinen and Champion, 1993a).

in situ and are not altered during heating in the laboratory. This is the basis of theoriginal method developed by Thellier and Thellier (1959) – the intensity of theancient-field being obtained from the slope of the NRM versusp TRM line (anArai (1963) plot).

In the classical Thellier/Thellier method of palaeointensity determinations, theNRM destroyed between successive temperature intervals (T1, T2 . . .) is comparedwith the TRM acquired in the same interval. This method has been modified by Coe(1967) in which the samples are given a series of double stepwise heatings. In thefirst, the partial NRM (p NRM) between two temperaturesT1 andT2 is removedby cooling in zero field, and, in the second, a partial TRM (p TRM) is acquiredover the same interval by cooling in some laboratory field. If no alteration in thesample has occurred in heating to higher temperatures, the ratiop NRM/p TRMshould be constant. i.e., the Arai plot should be a straight line (see Figure 1). In alater paper, Coe et al. (1975) have given a number of criteria that should be metbefore the Arai plot is accepted. The reliability of palaeointensity determinationsby this method has been argued by Aitken et al. and Walton in two back to backpapers in Reviews of Geophysics (26, 1–25, 1988).

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Viscous remanent magnetization (VRM) acquired at low temperature over timeperiods not reproducible in the laboratory can lead to a steeper slope in the Araiplot at lower temperatures. VRM will normally disappear at higher temperatures.“Back p-TRM checks” are carried out to monitor any chemical changes that mayhave occurred in the sample at high temperature by comparing ap TRM acquiredat one temperature with ap TRM acquired later at the same temperature, but afterthe sample had been subjected to a higher temperature demagnetization.

A further disadvantage of the Thellier/Thellier method is the length of timeinvolved because of the large number of heatings that are required. To overcomethis drawback, attempts have been made to use AF demagnetization methods,which only require one heating. Shaw (1974) suggested a method which wouldalso detect any alteration of the sample on heating. In his method, a comparison ismade of the AF demagnetization spectrum of anhysteretic remanent magnetization(ARM)∗ before and after heating. A plot of ARM 1 (before heating) versus ARM2 (after heating) for successive demagnetization fields will give a straight line ofunit slope if there has been no alteration. Kono (1978) and Rolph and Shaw (1985)have suggested a correction to the Shaw method which can be applied to thermallyaltered samples.

Haag et al. (1995) obtained absolute palaeointensities using the Thellier/Thelliermethod for a number of historic basaltic lavas from Mt. Etna which have similarchemical compositions, but different physical properties (e.g., porosity) because oftheir different cooling histories. The use of historic and recent lava flows meansthat the actual field intensity is known during the cooling time of the lava, sothat the validity of the Thellier/Thellier method can be tested. They found thatmineral changes and changes in the domain state at low temperatures can havea profound effect on the magnetic properties at high temperatures. Good agree-ment with the true field intensity was obtained if the individual low temperaturesegments of the Arai plot are used and the regression line includes the lowesttemperatures. On the other hand, using the Thellier/Thellier method with linear andreproducible behaviour only at high temperatures can give incorrect results. Pos-sible non-linearity of the Arai plot at low temperatures is often caused by viscousremanent magnetization (VRM). Non-ideal behaviour at high temperatures resultsfrom physico-chemical changes during repeated heatings. Walton et al. (1993) useddirect micro wave excitation of the magnetic grains to obtain a TRM without sig-nificantly heating the bulk sample. Thus mineral alteration owing to heating ofthe non-magnetic matrix is avoided. Using this technique with the conventionalThellier/Thellier method, Shaw et al. (1996) significantly reduced the scatter ofarchaeointensity measurements from Peruvian ceramics. In a later paper, Walton etal. (1996) used the method to obtain the intensity from samples of Egyptian pottery.

∗ ARM is the magnetization installed in samples by the application of a strong alternating fieldwhich is slowly reduced to zero in the presence of a weak direct bias field.


Valet et al. (1996) investigated the possibility of correcting absolute palaeoin-tensity experiments for changes in magnetization caused by magnetomineralogicalchanges during heating. They claim that such changes can be assessed by thep TRM checks. Their method involves additional, detailed analyses includingp

TRM checks at each temperature step and repeated measurements at the sametemperatures. Their method also depends on a knowledge of both the intensity anddirection of the field. They found that magnetomineralogical changes usually occurfor grains with blocking temperatures lower than the last heating step. They appliedtheir method to recent and historic lava flows from different areas and increased therecovery of reliable absolute palaeointensity determinations by about 50 per centwhen heating was carried out in vacuum and up to 65 per cent when heated in air.

McClelland and Briden (1996) have shown that failure ofp TRM checks doesnot necessarily indicate that alteration of high-blocking temperature material hasoccurred. They proposed a new experimental sequence for Thellier/Thellier typepalaeointensity determinations which allows monitoring and correction of alter-ation. They point out that when alteration begins at low temperature and continuesthroughout the experiment (which then fails in standard analyses), the resultingunblocking temperature (Tub) spectrum of the alteration product does not neces-sarily overlap with the wholeTub spectrum of the NRM, becauseTub depends onthe physical characteristics of the new grains, not on the temperature at which theywere formed. NRM may thus survive in an “uncontaminated” higherTub window.McClleland and Briden illustrate their new method on Palaeozoic igneous rocksfrom the Ntonya Ring Complex, an alkali-syenite intrusion in Malawi, in whichthey successfully separated TRM from multi-domain IRM despite large alterationduring laboratory heating.

Finally it must not be forgotten that only spot readings of the field can be ob-tained from lava flows and that it is impossible to determine their absolute ages– a sequence of lava flows may have been erupted over a few years or thousandsof years or more. In addition, palaeointensity measurements of volcanic rocks arelimited by the relative scarcity of volcanic eruptions on continents over the past100 ka.

In contrast to the spot readings provided by lava flows, sedimentary rocks givea much more continuous magnetic record. However there are difficulties in obtain-ing palaeointensities from sedimentary rocks, mainly because there is no soundtheoretical basis for the acquisition of remanence. We cannot produce in the labo-ratory the conditions under which the rocks acquired their remanence (depositionalremanent magnetization, DRM or post depositional magnetization, PDRM). Thefundamental assumption for sedimentary palaeointensity studies is that the NRMis linearly related to the magnetic field in which it was deposited. However theNRM will also depend on changes in magnetic concentration, grain size etc. Thusonly relative intensity changes can be determined by normalizing the NRM with asuitable factor to remove the influence of all variables except field intensity.

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The best parameter for normalization is still a matter of debate – isothermalremanent magnetization (IRM), low field magnetic susceptibility (χ), and anhys-teretic remanent magnetization (ARM) have all been used. Estimates of relativeintensity are then obtained by dividing the NRM by IRM,χ , or ARM. None ofthese estimates takes into account the possible contribution of viscous remanentmagnetization (VRM).

Saturation IRM was first used as a normalizing parameter by Johnson et al.(1948), and Nakajima and Kawai (1973) used it to estimate the secular variationfor the past 60 ka from sediments in Lake Biwa, Japan. Arguments in favour ofIRM have been given by Hartl et al. (1993). Harrison and Somayajulu (1966), intheir study of reversals of the Earth’s magnetic field, were among the first to useχ

as a normalizer. Advocates in favour of ARM include, amongst others, Johnson etal. (1975), Levi and Banerjee (1976), and King et al. (1983). Opdyke (1972) andOpdyke et al. (1973) advocate the use of IRM or ARM. Levi and Banerjee (1976)list the assumptions made in obtaining palaeointensities from sediments and Kinget al. (1983) give a critical evaluation of the hypothesis that normalized remanenceis a measure of relative palaeointensity and summarize the requirements for us-ing ARM, which, on balance, seems to be the favoured parameter. The relativemerits of these normalizers and other problems in determining palaeointensitiesfrom sedimentary rocks have been discussed in detail in an excellent review byTauxe (1993). In a later paper, Tauxe et al. (1995) developed a “pseudo Thel-lier/Thellier” method for determining palaeointensities which allows the removalof unwanted VRM improving agreement among different records and allowing anestimate of the uncertainties in the results. Using 19 samples from a core from theOntong Java Plateau, they showed that the three methods of normalization maybe systematically biased by unremoved VRM and that agreement between themdoes not guarantee a sufficiently rigorous test for reliability. In addition, as withpalaeointensity measurements from lava flows, there is the problem of determiningthe time when the sediments acquired their magnetization.

Apart from these experimental difficulties, there are other problems in obtainingreliable palaeointensity measurements. During periods of low intensity, it would beharder to align the magnetic grains – and this could vary with the palaeomagneticrecorder (TRM or DRM). In the case of magnetic grains in sediments, the physicalalignment is probably determined by hydrodynamic and gravitational forces, pre-venting the grains from being aligned by the magnetic field. Finally it should beemphasized that lava flows record the total field intensity i.e., both the dipolar andnon-dipolar components of the field. In contrast, the magnetization of sedimentstakes place over much longer time periods and marine sedimentary recorders arethus mostly sensitive to the dipole field intensity.

Constable and Tauxe (1996) suggested a different method of calibrating relativepalaeointensity records by sampling the field during a polarity transition whenthe contribution from the axial dipole is zero. The intensity record at the time ofthe reversal (when the inclination changes sign) then contains only the non-axial


dipole field (NAD). They assume that the NAD contribution to the field is invariantwith geographic position and, from absolute palaeointensity data, they estimate itsmagnitude to be 7.5µT. They then multiply the record by a constant factor to givethe transitional intensity a value of 7.5µT. In this way the whole palaeointensityrecord can be calibrated. With these assumptions, the average palaeointensity overeach record should give the variation expected from a dipole field as a function oflatitude. To test their method, Constable and Tauxe used relative palaeointensityestimates obtained by Hartl and Tauxe (1996) from 12 globally distributed pelagicsedimentary cores spanning the Matuyama-Brunhes boundary. They found thattheir method predicted the variation of the dipole field with latitude with the sameaccuracy as that obtained from absolute palaeointensity records.

3. Archaeomagnetic Results

McElhinny and Senanayake (1982) analysed all available archaeomagnetic inten-sity data for the past 50 ka. They list full references to earlier compilations andpapers, noting errors, omissions and duplication. Most of the data are from the past12 ka and this data set is strongly biased to regions lying between 0 and 90◦E inthe northern hemisphere (64% of the data). Figure 2 shows global mean dipolemoments with 95% errors for the past 12 ka and suggests a maximum around 500BC and a minimum around 4,500 BC. The mean dipole moment over the past 10 kais 8.75×1022 Am2. There are no data for the period 12–15 ka and for the period 15–50 ka only 14 values were available. Ten of these are very low-about 1/2 that overthe last 10 ka (∼ 4.44×1022 Am2). The other 4 intensities were very high and wereobtained from Lake Mungo, Australia at the time of a (possible) excursion of theEarth’s magnetic field. Due to averaging, the global intensity curve reflects changesin the main dipole and does not give information on the more rapidly varying non-dipole field. Regional archaeointensity studies are necessary to obtain informationon the strength and drift of the non-dipole field and to detect any periods of rapidintensity changes. Since we are more interested in changes in the dipole field, noattempt will be made to review the many analyses of archaeointensity data for thelast few ka. These analyses reveal differences in records from different parts ofthe world and reflect the changing non-dipole field. This is well illustrated in twostudies on data from China and Finland.

Shaw et al. (1995) obtained intensity variations of the Earth’s magnetic field forthe past 7.5 ka from ancient Chinese ceramics. Although major features of the fieldsin China are in general agreement with those of the global average field modelof McElhinny and Senanayake (1982), there are detailed differences. The authorssuggest that these differences may be the result of local non-dipole fields or lack ofresolution in the global model. It is interesting, however, that high palaeointensityvalues around 3.5 ka ago have been reported in China, Japan and Hawaii suggestingthat this feature may reflect a true maximum in the dipole field. There are more

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Figure 2.Global mean dipole moments with 95% errors for 500 year-intervals (to 2000 BC) and for1,000 year-intervals (prior to 2,000 BC) from archeomagnetic data. The numbers on the curve arethe number of points in each interval. (after McElhinny and Senanayake, 1982).

significant differences between the Chinese record and contemporaneous resultsfrom Greece (Aitken et al., 1989). Shaw et al. suggest that this could be the resultof the movement of a large non-dipole anomaly.

Constable (1985) obtained relative intensity estimates of the magnetic field fromcores from two volcanic crater lakes (Lake Barrine and Lake Eacham) both ofwhich are on the Atherton Table land of northern Queensland, Australia. ARMwas used as a normalizing paramter for the NRM. The chronology was basedon radiocarbon dating and goes back 6 ka for Lake Eacham and 14 ka for LakeBarrine, although the Lake Barrine chronology is poorly constrained. The withinlake agreement is in general quite good, although the coeval time span is rathershort. Absolute intensity measurements have been obtained by Barton |and Barbetti(1982) from archaeological work on aboriginal fireplaces in south-eastern Australiaspanning the last 7 ka. The agreement with the relative intensity estimates fromLake Eacham is good despite the rather large separation of the sites, showing ahigh around 3.5 ka ago following a low around 4 ka.

Pesonen et al. (1995) obtained intensity measurements of the Earth’s magneticfield in Finland for the past 6.4 ka using bricks, potsherds and baked clays. Theirresults show an increase in intensity from 4360 B.C. to a maximum at 500 AD(about twice the present value), after which it decreases to its present value. Thehigh at 500 AD is not seen in Bulgaria or Japan – the contemporaneous Bulgar-ian curve shows a broad minimum and the Japanese a broad plateau at that timerespectively.


Daly and Le Goff (1996) have compiled an updated world secular variation database using archaeomagnetic data from the last 2 ka. Values of the declination andinclination have been calculated at 25 year intervals from 9 world sites – Meri-den (England), Paris (France), Sofia (Bulgaria), Kiev (Ukraine), Gori (Caucasus),Kyoto (Japan), Arizona (USA), Arkansas (USA) and Meso-America (Central andsouthern Mexico, Guatemala, Honduras and Salvador). Intensity data were avail-able only from Bulgaria, the Ukraine, Caucasus and Japan. The curves clearly showthe regional differences in the more rapidly varying non-dipole field. Daly and LeGoff plan to give the data from volcanic rocks and lake sediments in a secondpaper. Kovacheva (1997) has given a compilation of all archaeomagnetic data fromBulgaria for the last 8 ka.

4. Intensity Measurements for the Last Few 100 ka.

There have been many estimates of palaeointensities of the geomagnetic field fromlava flows on Hawaii and these will be considered first. Most were from flows<5 ka old and very few from flows >10 ka – no attempt will be made to reviewthe early work. Hagstrum and Champion (1995) have compiled a palaeomagneticrecord of the geomagnetic PSV for the last 4.4 ka based on 191 historical and14Cdated lava flows on Hawaii. For earlier times, Mankinen and Champion (1993a)estimated the intensities of14C dated Hawaiian lava flows over the last 12 ka. Theirresults, together with earlier estimates and the global mean dipole variation curveof McElhinny and Senanayake (1982) are shown in Figure 3.

It appears that there has been little change in the Earth’s VDM between about12 and 5 ka B.P. and that the non-dipole field must have been virtually absentduring that time. On the other hand, the non-dipole field appears to have been quitestrong from about 5 ka to almost the present time. In a later paper, Mankinen andChampion (1993b) obtained more intensity measurements from Hawaii rangingin age from 31–13.5 ka BP. They estimated that the intensity of the geomagneticfield was reduced on average by∼35% between about 45 and 10 ka ago. Duringthis period of very low intensity, their Hawaiian data indicate that the field wasstronger between∼29–19 ka ago.

A more recent investigation of the geomagnetic field intensity in Hawaii hasbeen carried out by Garnier et al. (1996a), who obtained a record from 100 lavaflows going back to 42 ka from a core from the SOH-4 well. Assuming a linearextrusion rate between the present and the 42 ka date obtained with an unspikedK-Ar technique, this corresponds to an average of one flow every 420 a. Howevertheir results must be viewed with some caution. The palaeointensities were ob-tained from only one or two samples per lava flow which is insufficient to give areliable estimate. The value of the palaeointensity determined from one sample canbe very different from that obtained from other samples from the same flow, sincethe magnetic mineralogy is not uniform in a lava flow. Moreover their date of 42 ka

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Figure 3. Virtual dipole moments (VDMs) as a function of age for the Hawaiian14C-dated lavaflows. Dashed curve is a rough fit (by eye) of all paleointensity data from Hawaii. Global-meandipole variation curve (McElhinny and Senanyake, 1982) for Holocene time is shown with its 95%confidence envelope (shaded area). Dashed horizontal line is the worldwide average VDM for thepast 10,000 years (McElhinny and Senanayake, 1982). Circles are from Mankinen and Champion(1993a), triangles from Coe et al. (1978), and squares from Tanaka and Kono (1991). Labels are flowdesignations used by the authors. Open symbols are values based on fewer than three determinations.(after Mankinen and Champion, 1993a).

has a large estimated error of±15 ka. Thus, although their intensity record showssome agreement with other values obtained for Hawaii (Mankinen and Champion1993a, b; Tanaka and Kono, 1991), it is not surprising that there are differences –particularly with those obtained from other areas.

Garnier et al. (1996b) later obtained palaeointensity measurements for the last400 ka from cores drilled into the Mauna Loa and Mauna Kea volcanoes. Theresults obtained by the Thellier/Thellier method were based on only one sampleper flow. One14C date was obtained at∼40 ka and 9 K/Ar dates between∼105–400 ka – some of these later dates have large error bars. Thus the same criticismof their 1996a paper applies to their later paper – a fact which is recognized bythe authors themselves who say that “the preliminary results obtained here mustthen be examined with particular caution”. One feature of the records in both theirpapers which is of some interest if substantiated by further work, is the presence oflarge amplitude (20–25µT), rapid (4–6 ka) fluctuations of the intensity which arenot clearly seen in the earlier records from Hawaii.

Brassart et al. (1977) obtained absolute palaeointensities from 8 out of 10 lavaflows from Kohala Mountain, Hawaii with K/Ar dates ranging between 60 and 400ka. Using a new technique developed by Valet et al. (in press) to apply correctionsto thep TRMs, they were able to increase the success rate with 50 per cent addi-tional determinations. The correction was used for samples that showed changesin their ability to acquirep TRM during heating but did not show acquisition of


Figure 4.Comparison between the VADMs obtained by Brassart et al. (1997) and other results fromvolcanic and sedimentary records.4 = volcanic records (0–40 ka),# = Mt. Etna,2 = La ReunionIsland,5 = Big Island (Hawaii), = Brassart et al., o = synthesis of sedimentary records (Sint -200). After Brassart et al. (1997).

chemical remanence. Their results are in good agreement with other palaeointen-sity estimates obtained for the same period (see Figure 4) confirming the existenceof a significant decrease between 160–120 ka leading to a period of low intensity.The results are also consistent with the synthetic record of relative palaeointensityobtained for the past 200 ka from deep-sea sediment cores (Guyodo et al., 1996).Unfortunately detailed comparison with other records for the period 200–400 kaare not possible because of large error bars in age determinations which can reach100 ka.

Gonzalez et al. (1997) studied the secular variation in Central Mexico over thelast 30 ka. Samples (dated by14C) were taken from lava flows from 13 differentvolcanoes from two sites, both part of the Transmexican Volcanic Belt. Intensitieswere obtained from several samples from each site using both Thellier/Thellier andShaw methods. Variations of the VDM range from 3.1± 0.4 Am2 to 14.9± 2.3Am2 and show a gradual increase over the last 30 ka with a peak at∼9 ka in generalagreement with results obtained from both lavas and sedimentary rocks from otherparts of the world. The high VDM (14.9 Am2) at about 2 ka is similar to that foundby Tanaka and Kono (1991) from a Hawaiian lava dated at 1.8 ka.

Tric et al. (1992) obtained high resolution records of the relative palaeointensityof the geomagnetic field for the past 80 ka from 5 marine cores – 3 from theTyrrhenian Sea, one from the eastern Mediterranean and one from the southern

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Figure 5. Comparison of palaeointensities between sedimentary records for the last 80 ka andarcheomagnetic and volcanic data sets for the last 40 ka. (after Tric et al., 1992).

Indian Ocean. Rock magnetic analayses showed down core uniformity of the sed-iments in terms of magnetic minerology and grain size, and ARM, SIRM orχ

could be used to normalize the NRM. The results correlate well with archaeomag-netic data before 10 ka and with volcanic data for the last 40 ka (see Figure 5).Intercore correlations and chronological data were provided byδ18O data and bythe identification of tephra layers dated on land (Paterne et al., 1986). The dipolefield moment shows large scale changes with broad minima centered around 20 ka,and 38 ka where it fell to 22 per cent of its present value and around 63 ka where itfell to 28 per cent. These lows alternate with periods of higher intensity around 8,50 and 78 ka. Between 20 and 40 ka there are two medium amplitude oscillations.

Meynadier et al. (1992) obtained relative magnetic field intensities for the last140 ka from 3 marine cores in the Somali basin, Western Indian Ocean – a timeversus depth correlation was established from theδ18O record. ARM was used asa normalizing parameter for the NRM. The quasi-cyclic pattern for the past 80ka confirms the results obtained in the Mediterranean by Tric et al. (1992) and inaddition shows a broad change in the field intensity to lower values between 80and 140 ka, reaching a minimum value of one-third that of the present field, 115 kaago. It is interesting that this coincides with an excursion of the Earth’s magneticfield – the Blake event. Low field intensity was also observed during the Laschampevent about 40 ka ago.

Thouveny et al. (1993) obtained relative intensities of the Earth’s magneticfield over the last 70 ka from a sedimentary sequence in a maar crater lake, Lacdu Bouchet, Massif Central, France. The sedimentation rate (0.19–0.39 mm/yr) is


considerably higher than that in the Mediterranean cores (0.06–0.1 mm/yr) stud-ied by Tric et al. (1992). Dating was estimated in weighted average regressionlines computed between independent chronological data and ARM was used asthe normalizing parameter for the NRM. The results of Thouveny et al. showextremely low values of the geomagnetic field (7.7–15µT) between 45 and 32 kawith medium and high values after 30 ka. Between 30 and 10 ka there are severalshort period, large amplitude fluctuations which are probably associated with non-dipole field variations. Prior to 45 ka there are extra large amplitude, short periodoscillations with weak intensities between 70 and 60 ka.

It is interesting to compare the results of Thouveny et al. (1993) with thoseof Tric et al. (1992) (see Figure 6). In the interval of low intensity (45–30 ka),there are two minima at 33 and 43 ka in the Lac du Bouchet record, but only oneat around 39 ka in the Mediterranean record. The major difference, however, isin the large oscillations between 25 and 12 ka which accompany increasing fieldintensity in the Lac du Bouchet record, but with decreasing field intensity in theMediterranean record. Both investigators found a low in the intensity around 40ka ago, the time of the Laschamp event yet both could find no departures of thegeomagnetic vector from the normal direction – Lac du Bouchet is only 100 kmfrom the Laschamp site in La Chaine des Puys. Both investigators concluded thatthe duration of the Laschamp event could not have exceeded a few hundred years.Ohno et al. (1993) investigated the record of the magnetic field recorded over thepast 35 ka in a sediment core from off Shikoku, southwest Japan. Between 25 and12 ka oscillations in the intensity of the field were superimposed on an increasingtrend in agreement with the Lac du Bouchet record.

Roberts et al. (1994) analysed a 15 m succession of Middle/Late Pleistocene la-custrine sediments from Lake Chewaucan, southern Oregon in order to see whetherthey showed the same relative palaeointensity record as that seen in deep-sea sedi-ments. Normalization of the NRM by ARM, SIRM andχ gave essentially the sameresult, indicating that the effects of variations in magnetic mineral concentrationhave been removed. Ages for the youngest part of the sequence were obtainedfrom 14C and thermoluminesce (TL) dating of tephra layers and are believed tobe accurate to within∼10 per cent. TL and K/Ar dates for the lower part of thesequence have large standard errors, and Roberts et al. believe that the chronologyof Lake Chewaucan is sufficiently reliable only for the upper part of the sequencei.e., from∼65–105 ka. Reasonable agreement was found for this period with the re-sults of Tric et al. (1992) from the Mediterranean Sea and with those of Meynadieret al. (1992) from the Somali Basin, all three records showing a short intensitymaximum at∼70 ka and a broader, stronger maximum of∼80 ka (see Figure 7).The record from Lake Chewaucan is also in fair agreement with that of Meynadieret al. showing a broad intensity low from∼90–120 ka – the record of Tric et al.does not extend beyond∼80 ka.

Schneider (1993) used data from the topmost sections of two cores from SuluSea sediments to estimate geomagnetic intensity variations over the past 110 ka.

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Figure 6.Comparison of the relative palaeointensity record from Mediterranean cores top (Tric etal., 1992) with the record from Lac du Bouchet, bottom. (after Thouveny et al., 1993).

The chronology is based on published oxygen isotope and radiocarbon data. Theoxygen isotope age model agrees with the available14C data above 245 cm and wasused at greater depths. Schneider estimated the variation in the strength of the geo-magnetic field using low-field magnetic susceptibility as a normalization parameterfor the palaeointensity record calibrated against previous Holocene age estimatesof absolute intensity (McElhinny and Senanayake, 1982). Both cores show a rapid


Figure 7. (a) Sedimentary palaeointensity records from 60 to 110 ka from the Mediterranean Sea(Tric et al., 1992; bold line) and from the Somali Basin (Meynadier et al., 1992; thin line). (b) TheNRM/ARM record from the same interval for Lake Chewaucan (after Roberts et al., 1994).

decline in intensity between 50 and 40 ka. However there is a difference betweenthe records of the two cores between 40 and 30 ka – core 769B shows a partialrecovery in intensity during this interval which is not shown in core 769A.

The overall similarity of the Mediterranean (Mazaud et al., 1991, Tric et al.,1992) and Sulu Sea records is striking. Both regions show distinct low intensityintervals near 75–65 ka, near 40 ka and near 20 ka. Both also show a short, rel-atively low intensity interval near 15 ka – less marked in the Sulu Sea record.

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The consistency of both timing and character of the major palaeointensity shifts inthese widely separated basins suggests that the records do reflect significant dipolarvariations in geomagnetic intensity. The most significant difference between theMediterranean and Sulu Sea records lies in a 15 ka interval centered at approxi-mately 30 ka. The Mediterranean results show a double-peaked recovery of fieldintensity between the 40 and 20 ka lows, whereas the Sulu Sea records generallyshow low intensities from 40–20 ka. This differerence may reflect a genuine differ-ence in the relative geomagnetic field intensity experienced in these two regions.If this is the case, it would seem that the geomagnetic field may at times have hadlarge non-dipole components during this period of generally lowered field strength.

Mejia et al. (1996) used submarine basaltic glass to obtain palaeointensitiesusing samples for the last 69 ka, recovered from the northern East Pacific Rise.Dating was based on their distance from the ridge axis and U-series disequilibriameasurements. They obtained good agreement with the International GeomagneticReference Field for a sample erupted in the 1980s. Their results show a generalincrease in intensity from 69 ka to the present, though varying from as much as 50%higher than the present field to as low as 26%. The results are in broad agreementwith other absolute and relative palaeointensity records.

Schwartz et al. (1996) obtained relative palaeointensities from two cores fromsediments on the Blake Outer Ridge in the western North Atlantic Ocean coveringthe period 12–71 ka. Their records (see Figure 8) are remarkably similar to thosefrom Lac du Bouchet (Thouveney et al., 1993) and the Mediterranean Sea (Tric etal., 1992) – in particular low intensities near 40 ka and 60 ka. Dating was based onδ18O for the Blake Outer Ridge and Mediterrean records, and that for the Lac duBouchet record on14C and pollen stratigraphy. However, in spite of the good agree-ment between the three records, Schwartz et al. are concerned that their record issignificantly correlated with the down-core ratio of ARM/χ . This ratio is a measureof changes in the relative grain size of magnetite (the primary magnetic mineral inthe sediments) and is controlled only by the local depositional environment andindirectly by global climate. Their concern is that the European records may besimilarly biased by climatic and other environmental factors. They caution thatextreme care must be taken to remove any such magnetic influences from sedimentpalaeointensity records before they can be used as quantitative estimates of the pastintensity of the Earth’s magnetic field.

Peck et al. (1996) carried out palaeomagnetic studies on sediment cores fromthe Selenga prodelta region of Lake Baikal, Siberia. Detailed laboratory analysesshowed that the cores satisfied the criteria (Tauxe, 1993) for obtaining relativepalaeointensities and that the records reflect variations in the geomagnetic field andnot in the lithology. Accelerator mass-spectrometer (AMS)14C ages were obtainedfor the last 23 ka. Since Lake Baikal is undersaturated with respect to CaCo3, δ18Ocannot be used for dating and ages earlier than 23 ka ago, back to 84 ka, were deter-mined from rock magnetic climate proxy data. The relative palaeointensity recordfrom Lake Baikal correlates well with records from the Mediterrean Sea (Tric et


Figure 8. A comparison of relative palaeointensity records from the Blake Outer Ridge, Mediter-ranean Sea and Lac du Bouchet, along with the ratio ARM/χ (inverted). Although agreement amongthe three records is remarkably good, there is also a significant correlation with the rock magneticrecord e.g., dashed lines indicate the correlation between two intensity lows and ARM/χ peaks. (afterSchwartz et al., 1996).

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al., 1992) and with those from the Somali Basin (Meynadier et al., 1992) – seeFigure 9. The three records are from distinct sedimentary environments from twoseparate ocean basins 8000 km from Lake Baikal. However in some intervals theLake Baikal record is in better agreement with the record from the Mediterrean Sea,whilst in other intervals, it is in better agreement with that from the Somali Basin.Peck et al. believe that, by correlating continental and marine records, relativegeomagnetic intensity stratigraphy has a potential resolution of 7 ka.

Tric et al. (1994) extended the record of absolute palaeointensities back to 160ka using distinct lava flows from Mount Etna (Sicily). One aim of their studywas to improve and extend the calibration of relative palaeointensities obtainedfrom sedimentary sequences for the period 60–160 ka. Precise ages were obtainedby K-Ar dating. Good agreement was found between their new results and thesedimentary record published for the same period (see Figure 10). For the period15–50 ka, a VADM of 4.3±1.5×1022 Am2 obtained from the Mount Etna samplescompares well with the mean value of 5.1± 1.1× 1022 Am2 obtained from thesedimentary record. For the period 0–160 ka, the VADM values from Mount Etnaand the sediments are 6.3±2.6×1022 Am2 and 5.2±0.8×1022 Am2 respectively.Their results indicate, as has been found in other studies, an appreciable decreasein intensity between 80–150 ka, centered at 115 ka (the Blake event?)

Rais et al. (1996) obtained palaeointensities using the Thellier/Thellier methodfrom a sequence of 70 successive lava flows from the Piton des Neiges volcano onthe island of La Réunion in the Indian Ocean. Radiometric dating gives the age ofthe flows between 130 and 72± 3 ka. The intensity of the field varied between 13and 65µT, the average value 42µT being slightly higher than the present day fieldat La Réunion. A major characteristic of the field is the large scale intensity fluctu-ations with durations of the order of 10–30 ka (see Figure 11). Rais et al.’s resultsare consistent with those of Chauvin et al. (1991) at La Réunion for the period82–98 ka with a broad low around 95 ka. Another intensity low occurs between115 and 120 ka. On the other hand their results are significantly higher than thoseobtained from Mount Etna by Tric et al. (1994). Such large differences probablyreflect non-dipole components present during this period from two geographicallyseparated sites. The low at∼115 ka is about the time of the Blake event. If this isconfirmed, the record from La Réunion would be the first volcanic record of theBlake event from the southern hemisphere and would support the existence andglobal extent of this event.

Laj et al. (1997) obtained palaeointensities from 5 volcanic sections on theisland of Vulcano (Aeolian Islands, Sicily). Radiometric dates were obtained withan unspiked K-Ar technique and indicate that the sections span the interval 15± 2ka to 135± 4 ka with large gaps in between. Unfortunately only 40 samples outof 256 from 19 out of 70 flows yielded reliable palaeointensity determinations.Their results are roughly consistent with those from Mt. Etna (Tric et al., 1994)which is only 100 km away from Vulcano. The combined results indicate thatthe intensity of the geomagnetic field was significantly weaker than today in the


Figure 9. Comparison of relative palaeointensity records. (a) Lake Baikal plotted using AMS ra-diocarbon ages and the climate proxy age model. (b) Mediterranean Sea (shaded line) (Tric et al.,1992), Somali Basin (solid line) (Meynadier et al., 1992), and N.E. China (dotted line) (Yang et al.,1993) absolute intensity. (c) Lake Baikal relative palaeointensity record plotted after correlating tothe Somali Basin record. (after Peck et al., 1996).

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Figure 10.Virtual axial dipole moments (VADM) as a function of age. (after Tric et al., 1994).

Figure 11.Changes in the geomagnetic field intensity at La Reunion in the interval 130–72 ka. Thehorizontal line corresponds to the present value. Full circle = means with standard deviation (severaldata per flow). Open circle = single data per flow with standard error. (after Rais et al., 1996).


Mediterranean area around 110 ka. Absolute palaeointensity data are rather sparsefor the later period 90 to 135 ka, but comparisons can be made with results fromvolcanic sequences from the island of La Réunion in the Indian Ocean (Chauvinet al., 1991; Rais et al., 1996). Although the absolute intensity values for Vulcanoare in reasonable agreement with relative intensities obtained from sedimentaryrecords (calibrated with the volcanic data for the last 50 ka), the VADM valuesfrom La Réunion are consistently higher than those from Vulcano/Mt. Etna by afactor of 3–4. The reason for this large difference is uncertain, but Laj et al. suggestthat large, long-lived non-dipole components of the geomagnetic field may havebeen present at La Reunion between 90–130 ka.

Zheng et al. (1995) carried out a detailed analysis of the Earth’s magnetic fieldfor the last interglacial recorded in a late Quaternary loess-palaeosol sequence incentral China – the Hauanxian section in the northern part of the Loess Plateau, 300km northwest of Xian. Chronology was based on a comparison of the magneticsusceptibility record with deep-sea oxygen isotope stratigraphy. They obtained arelative palaeointensity record using the ratio of NRM to ARM, both thermallydemagnetized at 300◦C, and compared it with a stacked record from the Somaliand Mediterranean Seas (Meynadier et al., 1992, Tric et al., 1992) for the period70–138 ka. Although the non-dipole field will influence the two records differently,both show a trend of dipole decay and recovery with almost identical amplitude.

Stoner et al. (1995) obtained relative palaeointensity values for the Late Pleis-tocene from 3 deep Labrador Sea cores. Detailed studies showed that SIRM wasthe best normalizer for the NRM. The results are of particular interest since, un-like other marine palaeointensity records, these cores from high-latitude locationsare strongly affected by ice-sheet-ocean interaction and thus experience variablesedimentation rates and depositional mechanisms. A prominent feature in all threerecords is a palaeointensity low at∼31 ka. Above this low, there is a relative high,while below it, a prominent high (at 35.5 ka) is observed. Another prominent lowoccurs within isotope stage 4, below which is another interval of higher intensity.Although there is a strong correlation between the Labrador intensities and thosefrom a combined Mediterranean-Indian Ocean record (Meynadier et al., 1992, Tricet al., 1992), there are considerable discrepancies in chronology – e.g., the promi-nent intensity low at∼31 ka in the Labrador Sea record is seen at∼39 ka in theMediterranean, at∼37 ka in the Indian Ocean and at∼42 ka in the North Atlantic.The Labrador Sea chronology was based on AMS14C dates and planktonicδ18Ostratigraphy. The authors point out that in high latitude North Atlantic cores, theplanktonic oxygen isotope stratigraphy is difficult to interpret. In addition, becauseof ambiguities in isotopic records and rapidly changing sedimentation rates, agecontrol is relatively poor below the level of AMS14C dates. They thus confinetheir chronology to∼ the last 100 ka with much higher confidence and resolutionin the upper 40 ka of the record.

In a study of the long term secular variation over the last 200 ka, Yamazakiand Ioka (1994) examined 5 cores from the West Caroline Basin in the western

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equatorial Pacific. ARM was used as a normalizing parameter for the NRM. Datingwas by oxygen isotope ratios. Sharp intensity drops were found at∼40 and 190 kawhere the intensity was∼20% or less than the mean for the last 10 ka. A broaderlow was found at 100–110 ka. The first two lows were accompanied by directionalchanges in the field, suggestive of a short reversal event – the Laschamp and Biwa I(Jamaica) excursions. The intensity low at∼110 ka is close to the Blake event, butno intermediate directions were observed. Their results are in general agreementwith the data for the last 80–140 ka from the Somali Basin, the Mediterranean andthe Sulu Sea.

Yamazaki and Ioka (1994) also obtained a record of the inclination which wassuggestive of a long term secular variation with a period of 40–50 ka and an am-plitude of several degrees. This period is longer than the electrical diffusion timein the Earth’s core (∼15 ka or less), and they suggest that the cause of the longterm secular variation may be an external force which affects the fluid motion inthe outer core through core-mantle interaction. They note that the period is close tothat of one of the Milankovitch orbital functions – the obliquity (41 ka) and suggestthat there may be a relationship between the geomagnetic field, the Milankovitchorbital functions and climate. Such a link has been considered by a number ofauthors in the past. Meynadier et al. (1992) carried out a power spectrum analysisof the magnetic field intensity they had obtained from cores for the last 140 kafrom the Somali Basin and found two dominant peaks centred at 100 ka and 22–25ka and two smaller peaks at 19 and 43 ka. A longer time series is needed to cometo any conclusions on the significance of these spectral peaks. Meynadier et al.,however, point out that these values correspond to the frequencies of the Earth’sorbital parameters.

Guyodo and Valet (1996) combined 18 records of relative palaeointensities frommarine sediments to construct a synthetic curve for the last 200 ka, designatedSINT 200 – see Figure 12. Their major problems were inaccuracies and resolutionof the dating, which was based on oxygen isotopes. The problem was compoundedby the fact that different time scales had been used by different authors. Howeverall the records correlate fairly well when plotted on the same time scale. Onlythose records which satisfied the experimental criteria for relative palaeointensitydeterminations (Tauxe, 1993) were chosen and, in many cases, several parametersthat were used for normalization gave identical results. The curve SINT 200 wasobtained from the arithmetic mean of 17 individual records. Although the coresare not uniformly distributed, they are relatively wide spread around the Earth. Thedominant features of the synthetic SINT 200 curve reinforce previous observationsof low intensities at 40, 100–110 and 190 ka. Guoydo and Valet feel that it can beused as a reference for correlating and dating geomagnetic features longer than 10ka.

Weeks et al. (1995) obtained relative palaeointensities for the last 240 ka from4 cores from the central North Atlantic Ocean. Both ARM and IRM were usedas normalising factors for the NRM. Dating was from oxygen isotope ratios. Their


Figure 12.(a) Synthesis of relative palaeointensity records for the last 200 ka, referred to as SINT200. (b) The same as (a) after optimal adjustment of every individual record with SINT 200, with-out violating the isotopic correlations. (c) Arithmetic mean of the 15 most coherent records. (afterGuyodo and Valet, 1996).

results show broad similarities with those obtained by other authors from sedimentsdeposited in different environmental conditions in various locations around theworld – a low at 42 ka, a broad low in the interval 90–130 ka and another lowat 190 ka. A common signal would preclude climate variations as a major factor inthe relative intensity signal. Visual examination of their record does not show anyperiodicity or stationarity over the observed interval. The authors then carried out

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a spectral analysis. Although the first 80 ka suggests some oscillatory behaviour(period∼17 ka), the rest of the record does not support any claim for periodicity.

Schneider and Mello (1996) extended the palaeointensity record back to theMatuyama-Brunhes boundary using 4 cores from two sites in the Sulu Sea with ahigh sedimentation rate (∼10 cm/ka). Normalization by ARM, IRM andχ gavesimilar results. AMS14C dates were used for the first part of the record. For earliertimes, 11 astronomically calibrated ages from tuned isotopic records were chosen.For the last 10 ka, their results (Figure 13) are in good agreement with archaeo-magnetic determinations of the VADM (McElhinny and Senanyake, 1982) and forthe rest of the last glacial cycle they are very similar to earlier studies of marinesediments. In particular there is a sudden drop in intensity at∼46 ka followedby ∼20 ka of low intensities, more pronounced and prolonged in the Sulu Searecord than in the earlier studies. Their results are in closer agreement with thoseof Yamazaki and Ioka (1994) who found a low from 45–30 ka in cores from theCaroline Basin in the western equatorial Pacific than with those of Tric et al. (1992)and Meynadier et al. (1992) who found a recovery from the low at∼35 ka followedby a further fall between 25 and 15 ka in cores from the Mediterranean and SomaliBasin.

For the Brunhes chronology Schneider and Mello’s results (Figure 14) are infairly good agreement with those of Valet and Meynadier (1993) from the equator-ial Pacific, both studies showing relatively high intensities in the early, middle andlate parts of the Brunhes chronology with longer periods of low intensities in themid-Brunhes. Many shorter low intensity intervals (near 40, 200, and∼650–700ka) can also be seen in the two records. However there are differences in detail e.g.,the interval of relatively high intensity following the Matuyama-Brunhes transitionhas much higher amplitude in the eastern equatorial Pacific record than in that fromthe Sulu Sea, whilst the reverse is the case for the recovery of the field followingthe low intensity interval at∼200 ka.

Tauxe and Shackleton (1994) used relatively long (∼700 ka), isotopically dated(δ18O) pelagic carbonate sequences in a core from the Ontong–Java Plateau toinvestigate a possible link between the Earth’s magnetic field and its orbit. Satura-tion IRM was used as a normalizing parameter for the NRM. Spectral analysis ofthe relative palaeointensity record suggested periodic behaviour with a dominantperiod between 30 and 40 ka – which lies between the orbital functions of obliquityand precession. Peaks in the analysis thus appear not to be coherent with variationsin magnetic properties controlled directly or indirectly by climate, but are morelikely to represent true variations of the geomagnetic field. The authors concludethat the Earth’s orbit plays no part in the modulations of the Earth’s magnetic field.Some correspondence with eccentricity was seen, but the record is too short forany meaningful conclusions to be drawn.

Lehman et al. (1996) obtained relative palaeointensities of the geomagnetic fieldover the last 280 ka from 3 marine cores in the Açores area of the North AtlanticOcean which is further south than the cores studied by Weeks et al. (1995). The


Figure 13.Composite record of virtual axial dipole moments (VADM) for the last glacial cycle.Solid points are indicated where mean intensity could be derived from 2 or more holes. Error barsshow 2-σ standard errors of the mean. Shaded region indicates where only 2 estimates were available(i.e., where no formal errors could be determined), showing the range of those two estimates. Inset:calibration of Sulu Sea paleointensity results to the archeomagnetic compilation of McElhinny andSenanayake (1982) by scaling the Sulu Sea estimate to achieve the best fit. (after Schneider andMello, 1996).

cores were dated by oxygen-isotope data. Different normalising parameters gaveessentially the same results. Peaks around 18 ka and 23 ka, which reflect preces-sional changes in the Earth’s orbit, were found in the power spectra of NRM,χ ,ARM and SIRM. However these peaks were effectively removed by the normal-ization process and the authors believe that these frequencies seen in the NRM arean environmental feature and are not related to variations in the Earth’s magneticfield. Their interpretation is in disagreement with Meynadier et al. (1992) who sawa 23 ka component in a record from the Somali Basin which they considered to bea true geomagnetic signal. Many of the features in Lehman et al.’s North Atlanticrecord are in agreement with the main characteristics seen in other sedimentary andvolcanic records – e.g., low intensities around 40, 120 and 190 ka and highs at 50and 80 ka.

Schnepp (1996) has obtained palaeointensities for 18 rock units from the Qua-ternary East Eifel volcanic field (Germany) using the Thellier/Thellier method.Dating was by the40Ar/39Ar method and covers volcanic activity from 100 to 11ka ago. Mean palaeointensities range from 62–29µT which is within the range

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Figure 14.Top: Brunhes chron palaeointensity results obtained from deep-sea sediments from theeastern equatorial Pacific by Valet and Meynadier (1993). Bottom: Estimate of VADM for theBrunhes chron obtained from sediments from the Sulu Sea (Schneider and Mello, 1996).

expected for the secular variation (the present-day field strength is 48µT). Asurprising result is that the palaeointensity values are grouped around 50 and 30µTwith a gap at 40µT. The distribution is thus bimodal rather than Gaussian. In-tensities from the neighbouring West Eifel volcanic field show the same bimodaldistribution. However the global data set of 128 VDMs for the Brunhes chron(rocks >30 ka old) does not show this feature, if the 14 transitional field values areexcluded. Thus the bimodal distribution found by Schnepp for the Eifel volcanicsdoes not seem to be representative of the Earth’s magnetic field during the Brunheschron.

The production rate of cosmogenic radioisotopes (such as14C and10Be) in theEarth’s atmosphere depends on the incident primary cosmic ray flux, solar activityand shielding by the geomagnetic field. Using the estimate by Tric et al. (1992)of the strength of the geomagnetic field over the last 80 ka, Mazaud et al. (1991)found that the geomagnetic field has been the main factor governing the productionof cosmogenic14C. This has caused problems with14C dating and corrections tothe time scale have to be made e.g., low values of the geomagnetic dipole momentin the interval 18–40 ka introduce a shift of∼2–3 ka in14C ages towards youngervalues (Mazaud et al., 1991).

The influence of the geomagnetic field on10Be production has been more con-troversial. Increases in10Be deposition during the last glacial maximum in thePacific and peaks in10Be concentrations in ice cores and the Mediterranean havebeen suggested to be due to a reduction in the intensity of the geomagnetic field.


Robinson et al. (1995) found a significant inverse correlation between the profilesof 10Be and the intensity of the geomagnetic field in a 80 ka core from the northeastAtlantic Ocean – in particular a rapid decrease in field intensity between∼40–50 kais accompanied by a rapid increase in10Be/9Be. They suggested that measurementsof 10Be/9Be in sediment cores might be used to give an estimate of palaeointensi-ties. Frank et al. (1997) followed up this suggestion and compared the geomagneticfield strength over the last 200 ka computed from a global stacked record of10Bedeposition in marine sediments with the field strength obtained from a stackedpalaeomagnetic record (Guyodo and Valet, 1996) (see Figure 15). It can be seenthat there is fairly good agreement between the two independently derived records– the only real difference occurring between∼115–125 ka. It is interesting thatRaisbeck et al. (1994) found no evidence in10Be records to support the saw-toothstructure of palaeointensity between reversals proposed by Valet and Meynadier(1993) – see Section 5.

5. Asymmetric Saw-Tooth Pattern of Palaeomagnetic Intensities

An interesting pattern of geomagnetic field intensities over the last 4 Ma was ob-served by Valet and Meynadier (1993) in their analysis of cores from the equatorialPacific Ocean. Relative palaeointensities were converted into VADMs by matchingthe data with the synthetic record of Meynadier et al. (1992) for the past 140 kawhich had been calibrated to volcanic VADMs for that period by Tric et al. (1992).The most striking feature in their record of intensities is an asymmetrical saw-toothpattern associated with reversals of the geomagnetic field – a gradual overall slowdecrease in intensity before a reversal accompanied with higher frequency, smallervariations followed by a very rapid recovery immediately following the directionalchanges in the field (see Figure 16). It should be noted, however, that during thelast 2 Ma there are a number of intensity drops which are not associated with fieldreversals or known excursions. Moreover, although there was a major field recoveryfollowing the last reversal, the Brunhes chron does not show a regular decrease offield intensity since the last reversal.

Even more surprising is their proposed relationship between the amplitude ofthe intensity recovery following a reversal and the duration of the subsequentpolarity interval, implying that the stability of a polarity interval depends on theamplitude of the intensity jump which followed the previous polarity transitioni.e., that the core has some memory of past reversals. Some support for this con-clusion has been given by Olson and Hagee (1990) on theoretical grounds. Theyshowed that small changes inRα, the magnetic Reynolds number based on theα effect, equivalent to small changes in the intensity or structure of outer coreconvection, can result in an irregular pattern of reversals. Their calculations predicta correlation between field strength and length of polarity intervals.

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Figure 15. (a) Global stacked10Be deposition rate which renders relative10Be production ratevariations as function of age. The dark grey areas in (a) and (b) mark the standard error of themean. The light grey bars mark periods in which the data show that the production rate of10Bewas increased and the field intensity was reduced. (b) Relative geomagnetic field intensity variationsderived from10Be production rate changes as function of age. The values between 140 and 205 kyrrepresent 3-point running means. (c) Global stacked record of field intensity variations based on astacked paleomagnetic record (Guyudo et al., 1996). The dark grey area marks 1 standard deviation(after Frank et al., 1997).

Meynadier et al. (1994) later obtained relative palaeointensities for the sameperiod (4 Ma) from marine sediments in the equatorial Indian Ocean, more than30,000 km away from the cores they recovered in the equatorial Pacific in theirearlier study. Different and independent methods yielded a time scale with a pre-cision better than 20 ka. Good agreement was also obtained between differentnormalization parameters. They found the same asymmetric saw-tooth pattern thatthey had seen in the Pacific Ocean to be also present in the record from the IndianOcean (Figure 17). Not only was the same pattern seen across every reversal, butthe short term fluctuations superimposed on the slow intensity decrease precedingthe reversal was also observed at both sites. The global character of these featuressuggests that they reflect changes in the intensity of the dipole field. The onlysignificant difference is for the time interval following the Upper Olduvai where


Figure 16.Relative variations of the dipole field intensity during the past 4 Ma. The polarity intervalsare indicated by horizontal bars (black/white shows normal/reverse polarities) and the position of thereversals shown by solid arrows. The excursions and the short events observed in previous studiesare correlated with intensity minima of the present record. In order to have a detailed view of theBrunhes chron, the upper and lower figures are not plotted on the same horizontal scale. (after Valetand Meynadier, 1993).

an almost reverse saw-tooth pattern is seen in the record from the Pacific core. Thecorrelation between the amplitude of the intensity recovery following a reversaland the duration of the subsequent polarity interval noticed by Valet and Meynadier(1993) in the Pacific record was confirmed in the Indian Ocean data (Figure 18).Valet et al. (1994) later compared the pattern of relative palaeointensities acrossthe Matuyama/Brunhes boundary as seen in sediments from the Atlantic, Indianand Pacific Oceans. Their results further strengthened the world-wide character ofthe asymmetric saw-tooth pattern – a slow intensity decay before the reversal anda very rapid recovery immediately following the transition.

Kent and Schneider (1995) also obtained relative palaeointensities across theMatuyama/Brunhes transition from two high resolution (8–11 cm/ka) sites in deep-sea sediments from the western equatorial Pacific. They found the record to be far

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Figure 17.Estimates of relative paleointensity for the last 4 Ma from the Pacific and Indian Oceans.Despite the very low resolution at the Indian Ocean site before 2.5 Ma, the two records are in verygood agreement. To enable plotting on the same vertical scale, the mean value of each record isnormalized to unity. (after Meynadier et al., 1994).

more complex than that from sediments with lower sedimentation rates. Two suc-cessive marked decreases in intensity were seen, each of about 15 ka wavelength.The full reversal in polarity directions is associated with the uppermost palaeoin-tensity decrease and took only∼2 ka. The authors believe the lower palaeointen-sity decrease is an independent geomagnetic feature that preceded the Matuyama/Brunhes transition.

Yamazaki et al. (1995) obtained relative palaeointensities for the Brunhes chronfrom 4 cores taken from different regions in low latitudes in the Pacific Ocean. Agecontrol was from oxygen isotope ratios and astronomical calibration. Their resultsare in good agreement with previous determinations. High frequency variations aresomewhat smoothed out because of the low sedimentation rate (∼10 m/Ma). Theasymmetric saw-tooth pattern of intensity variations across the Matuyama/Brunhestransition, first noticed by Valet and Meynadier (1993), was confirmed. Both recordsalso show an additional asymmetric pattern in mid-Brunhes – a quick growth from500–400 ka followed by a gradual decrease. The only discrepancy between the tworecords is a prominent intensity peak at∼650 ka in Valet and Meynadier’s recordwhich is absent in that of Yamazaki et al.


Figure 18.Relationship between the field intensity jump following every reversal and the durationof the subsequent polarity interval. The calibration in terms of VADMs relies on the compilation ofvolcanic data. (after Meynadier et al., 1994).

Thibal et al. (1995) confirmed an earlier study of Pozzi et al. (1993) that aprecise magnetostratigraphy can be obtained from continuous downhole measure-ments. Their results, from a site in the North Pacific, showed excellent agreementwith the geomagnetic polarity timescale of Cande and Kent (1992) down to a depthof 450 m below sea floor corresponding to an age of more than 8 Ma. Thibal et al.further showed that a continuous palaeointensity record can be obtained from asingle logging of a deep-sea sedimentary sequence. Figure 19 shows their relativepalaeointensities for the period 4.7–2.7 Ma which are in good agreement with, andextends by 0.7 Ma, the results of Valet and Meynadier (1993). Thibal et al. confirmthe asymmetric saw-tooth pattern of field intensity, observed by Valet and Mey-nadier. It is particularly interesting that Valet and Meynadier’s record was obtainedat low equatorial latitudes whereas Thibal et al.’s was from a high latitude. Valetand Meynadier had suggested that the length of a polarity interval is proportionalto the field intensity jump across the transition. Thibal et al. proposed a differentrelationship viz that the duration of a polarity interval is inversely proportionalto the mean rate of decrease in field intensity during that interval. More data arerequired to test the validity of these suggestions.

Laj et al. (1996a) obtained relative palaeointensities from two Late Miocenesections at Potamida and Kotsiani in Western Crete. The sections record two rever-sals bounding a normal polarity zone just below the Tortonian-Messinian boundary.

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Figure 19.Relative variation of the geomagnetic field intensity over the 4.7–2.7 Ma interval obtainedfrom down hole magnetic records at ODP site 884. The graph is shown in a polarized form, positivevalues for normal (black) and negative values for reverse (white) polarity intervals. Black circles areNRM/ARM ratios measured on specimens after AF demagnetization up to 20 mT. (after Thibal etal., 1995).

The astronomical polarity time scale for the Miocene (Hilgen et al., 1995) wasadopted. Normalization with ARM, SIRM andχ gave very similar results. Theauthors also carried out a spectral analysis of the NRM, ARM,χ and NRM/ARMrecords and showed that they are not related to the astronomical climate (Mi-lankovich) parameters. They thus feel confident that the records truly reflectchanges in the intensity of the geomagnetic field. The main feature of the two in-tensity records is the presence of large amplitude rapid changes which tend to maskthe longer term trends – this feature is also seen in records from other areas. Theauthors find no evidence for the asymmetric saw-tooth pattern seen in sedimentsby Valet and Meynadier (1993) for the last 4 Ma. Instead the first R→N transitionis characterized by a broad low in the intensity with no significant difference in itsvalue immediately before and immediately after the polarity change. The recoveryof the field is progressive and reaches a maximum more or less in the middle ofthe normal polarity zone. The field then progressively decreases until the next (N→ R) transition. Again no sudden recovery of the field is seen across this secondreversal.

Laj et al. (1996b) also carried out a new analysis of core KK78030 from the cen-tral equatorial Pacific. Mineral and magnetic analyses showed that the sedimentsmeet the criteria for reliable relative palaeointensity determinations and normal-


ization with ARM, IRM, andχ gave similar results. The core spans the last 1.8Ma and records the last 4 reversals. They found that the field intensity shows largeamplitude, fine scale fluctuations and gives no support to the saw tooth pattern.Verosub et al. (1996) also reanalysed data from core KK78030 but only from asmall section spanning the Jaramillo subchron and the Matuyama/Brunhes bound-ary. Like other palaeointensity determinations, Verosub et al. found high amplitude,rapid fluctuations superposed on longer term trends. They also saw some of thefeatures of the asymmetric saw-tooth pattern – a general decrease in intensity dur-ing the Jaramillo subchron and a very rapid rise following the Matuyama/Brunhestransition. However, as Laj et al. (1996b) found, when the whole core is considered,long term trends in the record can be seen which are lost in the shorter section. Thepartial support for the saw-tooth pattern seen by Verosub et al. (1996) in a shortsection of the core cannot be used to decide whether it is a common characteristicof polarity transitions.

Kok and Tauxe (1996a) have shown that unremoved long term viscous remanentmagnetization (VRM) can explain the saw-tooth pattern seen in some palaeointen-sity records for the last 4 Ma. They developed a model of VRM in a reversingfield and calculated the cumulative magnetization after a number of reversals. Acomparison of their synthetic cumulative remanance model is in fair agreementwith the pattern of relative palaeointensity observed in sedimentary sequences fromthe equatorial Pacific (see Figure 20). The authors are at pains to point out thatthough they have not proved that the asymmetric pattern is caused by long termVRM, nevertheless the success of their relatively simple model is encouraging.

In a later paper, Kok and Tauxe (1996b) re-examined the Gilbert/Gauss tran-sition. Thermal demagnetization caused the large post-transitional recovery of thepalaeofield to completely disappear. They were able to show that their cumulativeviscous remanence model (Kok and Tauxe, 1996a) gives an NRM at present that isvirtually the same as the previous results of Valet and Meynadier (1993), confirm-ing their earlier suggestion that the saw-tooth pattern seen by Valet and Meynadieris not of geomagnetic origin. It is interesting to note that Verosub et al. (1996) hadearlier maintained that the correlation of the high amplitude, fine scale featuresof their record with those of Valet and Meynadier rules out the possibility thatVRM is the explanation of the saw-tooth pattern as proposed by Kok and Tauxe(1996a, 1996b), claiming that if VRM were the explanation, it would erase or atleast obscure the correlation. Meynadier et al. (1997) have carried out a theoreticaland experimental study of the cumulative viscosity processes proposed by Kok andTauxe (1996a) and concluded that it cannot be responsible for the asymmetric sawtooth pattern.

Mazaud (1996) has developed an alternative model to explain the asymmetricsaw-tooth pattern seen in some intensity records from deep-sea cores. He showedthat the pattern can be produced if a large fraction of the NRM is acquired by thesediments at the time of deposition and the rest over a period of several hundredka after deposition (see Figure 21 for details). His model also does not damp out

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Figure 20.(a) NRM at present as a function of depositional age using the geomagnetic reversal timescale of Valet and Meynadier (1993). J. = Jaramillo; O. = Olduvai; K. = Kaena; and M. = Mammoth.(b) Model of a normal distribution ofτ with a mean of 50 Ma, and a standard deviation of 25 Ma(N = 500) (heavy line). Also shown are the sedimentary data of Valet and Meynadier normalized bytheir mean value. (c) Histogram ofτ values used.τ is a decay constant, depending on a number offactors such as saturation magnetization, grain volume, coercivity and temperature. (after Kok andTauxe, 1996).

the rapid intensity fluctuations which are observed superimposed on the saw-toothpattern. The difference between the models of Kok and Tauxe (1996a) and Mazaud(1996) is that in the latter model, the secondary magnetization is locked in phys-ically, whereas in the former model, it always relaxes to the varying equilibriummagnetizations.

Meynadier and Valet (1996) suggested that PDRM progressive reorientationsof magnetic grains might account for the asymmetric saw-tooth pattern. They car-


Figure 21.(a) The field input reproduces the geomagnetic polarity sequence (0–4 Ma) with a constantintensity during stable polarity periods. (b) The model assumes that a large fractionX of the mag-netic grains acquires its NRM at deposition time and thence remains unchanged through time. Theremaining fraction of the magnetic grainsY = 1− X acquires its magnetization progressively aftersediment deposition. The modelled intensity profile is obtained with a linear lock-in distribution overan interval1t after deposition for the fractionY of the magnetic grains.1t = 400 ka andX = 0.35.(c) Modelled intensity profile obtained with an exponential lock-in distribution for the fractionY ofthe magnetic grains characterized by a parameterτ = 200 ka andX = 0.35. (d) and (e) Comparisonwith the saw-tooth profile obtained by Meynadier et al. (1994). Dotted lines in the youngest few 100ka correspond to uncompletely blocked magnetization of the fractionY of the magnetic grains. (afterMazaud, 1996).

174 J. A. JACOBS

ried out a number of simulations of PDRM processes and concluded that such amechanism is unlikely. The only model compatible with the data requires that atleast 50% of the magnetization was acquired less than 15 cm below the surface ofthe sediment while the rest of the magnetization becomes locked-in progressivelybetween 10 and 30 m below the surface. Such a model is difficult to reconcile withpresent knowledge of compaction within the upper few tens of cm of sediment.

Tarduno and Wilkison (1996) investigated the possibility that palaeomagneticdata may be affected by non-steady magnetic mineral reduction leading to a “chem-ical lock-in” of magnetic particles. They define chemical lock-in as the level belowwhich biotic and abiotic chemical processes linked to redox reactions cease tocause appreciable changes in remanence. This concept is analogous to the phys-ical lock-in process in which increased sediment compaction eventually limits therotation and alignment of magnetic grains in the field. To check the feasibilityof this idea, they examined pelagic sediments on the Ontong Java Plateau in thewestern equatorial Pacific Ocean. The NRM data are suggestive of a chemical lock-in process in which some grains are formed in situ near and above the modernFe-redox boundary. The depth of this boundary suggests a delayed remanenceacquisition on time scales ranging from less than 40 ka to more than 425 ka. Thehigher values are compatible with delayed remanence acquisition models for thesaw-tooth pattern observed in some palaeointensity records associated with rever-sals of the geomagnetic field. It should be noted that the time scales of chemicallock-in are much greater than those associated with physical lock-in. Moreover thetime constant of VRM used by Kok and Tauxe (1996a) in their model is far longerthan previously thought in general.

6. Palaeointensities in Earlier Times

Tauxe and Hartl (1997) examined a 11 Ma record of the Earth’s magnetic field fromcores from the South Atlantic extending back in time an earlier analysis of coresfrom the same site by Hartl et al. (1993). An average sample spacing of 4 cm gaveone specimen per 4–8 ka. The many reversals seen in the records provided tie pointsto the marine magnetic anomaly time scale of Cande and Kent (1992, 1995). Nor-malization using ARM, IRM andχ gave consistent results. The overall shape of thepalaeointensity record seen in the earlier analysis of Hartl et al. is confirmed in thismuch longer record. The asymmetric saw-tooth pattern observed in some recordsis rarely seen by Tauxe and Hartl (1997) – rather their palaeointensity record showssteep-sided, roughly symmetric arches bounded by low intensity features varyingin time from∼40 ka to 10 ka. Tauxe and Hartl also carried out a spectral analysisof the palaeointensity record from 3 cores and found significant power with periodsin the range 30–50 ka. Although this is near the dominant period of the obliquity ofthe Earth’s orbit, they could not confirm any causal relationship between the two.Tauxe and Shackleton (1994) had earlier found a similar frequency in Brunhes-


age sediments from the Ontong Java Plateau, but were able to show that, in thatcase, the frequency of the geomagnetic “pulses” was unrelated to that of either theobliquity or the procession of the Earth’s orbit.

Prévot et al. (1990) compiled a list of all Triassic and younger intensity deter-minations of the Earth’s magnetic field based on Thellier/Thellier measurementsof volcanic rocks. They found large, long term changes in the strength of the field.From the K-T boundary to Pleistocene times, the average VDMs are close to thepresent value (8.75× 1022 Am2) for the last 10 ka, and they suggested that thestrength of the field has remained∼ approximately constant for the whole of theCenozoic. On the other hand, during most of the Mesozoic, the strength was onlyabout 1/3 of that during the Cenozoic (see Figure 22). From the lower Triassicto the lower Jurassic, the strength of the field again appears to be comparable tothat of the present field, the drop in intensity beginning∼185 Ma ago. The authorscaution, however, on the interpretation of their results in view of the sparseness andreliability of the data. Figure 22 also shows that for the period 0–160 Ma, changesin palaeointensity seem to correlate with changes in the amplitude of directionalpalaeosecular variation (PSV). Most of the Mesozoic dipole low also correspondsto a progressive decrease in reversal frequency.

Perrin et al. (1991) obtained palaeointensities, using Thellier/Thellier methods,for early Jurassic continental tholeiites from Britanny and the Iberian peninsula(Spain and Portugal). The results from these two dyke systems are essentially thesame (VDM∼3.5× 1022 Am2) confirming the existence of the Mesozoic dipolelow. Derder et al. (1989) on the other hand had earlier obtained palaeointensitiesnear present day values from the Newark Supergroup of eastern North America.Perrin et al. considered three possibilities to account for this discrepancy – errors inthe laboratory determinations of the intensities, significant age differences betweenthe tholeiites on opposite sides of the Atlantic Ocean and PSV. They were unable toresolve the differences, but still believe in the existence of a Mesozoic dipole low.The sparseness of data and a more exact chronology did not permit them, however,to give an age for the lower limit of the Mesozoic low.

Pick and Tauxe (1993) obtained high quality palaeointensity data from Thel-lier/Thellier experiments on recent and Cretaceous submarine basaltic glasses(SBG). The recent samples faithfully give today’s geomagnetic intensity at the site,giving confidence in using SBGs for intensity measurements – extrusion on theocean floor was actually observed so that both the age and the intensity of the fieldwere known exactly. Palaeointensities for the beginning and end of the CretaceousNormal Superchron (CNS) are only 45% and 25% respectively of today’s values.The data thus extend the Mesozoic dipole low into the CNS and do not support theexistence of a strong magnetic field during the Cretaceous. Pick and Tauxe (1993)suggested that the processes governing field reversals and long term changes inthe palaeointensity of the field are quite different. Prevot et al. (1990) had earliersuggested that the two processes are “probably decoupled”. McFadden and Merrill(1986) had earlier still argued that the physical processes that trigger reversals must

176 J. A. JACOBS

Figure 22. Geomagnetic virtual dipole moment, palaeosecular variation and reversal chronologysince the Triassic. Numbers by the dots are numbers of chronologic units. On the polarity timescale,normal periods are black and reversed periods white. (after Prevot et al., 1990).

come from an energy source that is independent of that which powers the mainmagnetic field.

There has been much controversy about the strength of the magnetic field dur-ing a superchron with different authors coming to opposite conclusions. Pal andRoberts (1988) and Larson and Olson (1990) maintain that the magnetic field isstrong during a superchron, arguing that a thin D′′ layer at the bottom of the mantlewould increase heat flow from the outer core (OC) to the mantle and so createmore vigorous convection in the OC. The associated heat loss would lead to astrong, stable, non-reversing field (a superchron). Loper |and McCartney (1986),McFadden and Merrill (1986) and Courtillot and Besse (1987), on the other hand,had earlier argued against a strong magnetic field during a superchron which they


claim represents a low energy state with infrequent instabilities correspondingto inactive periods in the D′′ layer when it is thick and heat transfer across thecore-mantle boundary relatively small. Both arguments seem reasonable, although,one wonders, that, if the average intensity was as low as Pick and Tauxe (1993)concluded for the entire CNS, why did not the field reverse? Prévot and Perrin(1992) later compiled a list of palaeointensity data obtained by the Thellier/Thelliermethod from magmatic rocks up to 3.5 Ga old. They found low values of the VDM(4–7× 1022 Am2) during the Permian which again does not support the claim ofPal and Roberts (1988) that the dipole moment was also appreciably greater duringthe Permian-Carboniferous Reversed Superchron (PCRS).

The debate still continues. Sherwood et al. (1993) obtained estimates of thestrength of the magnetic field during the CNS using lavas from Israel and India,with both the Thellier/Thellier and Shaw methods. Unfortunately the mineralogyof the rocks prevented good quality palaeointensity data being obtained by eithermethod. The mean VDM for the Israeli lavas was 4.7± 2.2× 1022 Am2) and forthe Rajmahal Traps in northeastern India, 5.5 ± 1.9 × 1022 Am2). Radiometricdating indicated that igneous activity occurred about the same time in both cases.Combining both data sets gives an estimated VDM of 5.4± 1.9× 1022 Am2) forthe beginning of the CNS. This is higher than the value obtained by Prévot et al.(1990) for the Mesozoic dipole low implying that the period of low dipole strengthhad ended before the onset of the CNS, contradicting the results of Pick and Tauxe(1993) who found that the Mezoic dipole low extended into the CNS.

Tanaka et al. (1995) analysed a global palaeointensity base from published datafrom volcanic rocks older than 0.03 Ma using both Thellier/Thellier and Shawmethods. Almost 90 per cent of the data are from the last 500 Ma. They found longterm variations of the Earth’s dipole moment with a broad minimum around 180–120 Ma in agreement with the earlier results of Prévot et al. (1990) (see Figure 23).However, as discussed earlier, this Mesozoic low cannot be definitely establishedbecause of insufficient site distribution and some contradictory data – more than 90per cent of the data for the period 120–210 Ma come from the Armenian region ofRussia and less than 10 per cent from Europe and North America. The analysis ofTanaka et al. (1995) also showed that a geomagnetic field existed more than 3000Ma ago with apparent lows at∼500 and 2000 Ma and an increase between 2500and 3000 Ma. However the data base is far too small to regard this as more thanspeculative.

Thomas et al. (1995) obtained 50 Thellier/Thellier measurements of the in-tensity of the Earth’s magnetic field from two large intrusive bodies of UpperCarboniferous age – the Great Whin Sill (GWS) of northern England and theMidland Valley Sill (MVS) of Central Scotland. Analysis of the direction of thefield indicates that samples from both sills lie within the PCRS, although they maynot be exactly contemporaneous. A mean palaeointensity of 22.9± 2.6µT (VDM5.9 × 1022 Am2) was estimated from the GWS – for the MVS the values weremuch lower (13.0± 0.5µT and 3.3× 1022 Am2). If the two bodies are not exactly

178 J. A. JACOBS

Figure 23.Long term variation of virtual dipole moments observed over the past 300 Ma. Broadminimum of the Earth’s dipole moment, the Mesozoic dipole low, preceded the Cretaceous normalsuperchron. (after Tanaka et al., 1995).

contemporaneous, such a difference (∼80%) is not unlikely. Over the last 12 ka,there is a similar (∼70%) variation in the VDM – 6.7× 1022− 11.3× 1022 Am2

(McElhinny and Senanayake, 1982). Figure 24 shows a summary of all VDMsfrom rocks younger than 360 Ma. For the PCRS, the Russian results indicate astrong dipole field, lending support for the model of Pal and Roberts (1988) andLarson and Olson (1991). However the remaining results from Northern Europeare inconclusive, indicating a dipole field supporting neither strong nor weak fieldmodels. Harcombe-Smee et al. (1996) later obtained palaeointensities from theMauchline lavas of SW Scotland which were erupted near the middle of the PCRS.The mean value of the intensity (13.6 ± 4.5 µT), corresponding to a VDM of3.2× 1022 Am2), is 40% of the present day value, and gives further evidence thatthe field strength was low during the PCRS.

Morimoto et al. (1997) carried out a palaeomagnetic study of a dolerite dykein West Greenland. K/Ar dating on plagioclase from the dyke gave a mean ageof 2750 Ma. Thellier/Thellier experiments on 12 samples gave a palaeointensityof 13.5± 4.4 µT, which, with the observed inclination (74.6◦), corresponds to aVDM of 1.9 ± 0.6 × 1022 Am2). This is about only one quarter of the presentvalue of the intensity. It must be stressed that this low value of the strength ofthe field is only a spot measurement since the width of the dyke (4 m) is too thinto record geomagnetic secular changes longer than about 10 yr. Low values ofthe palaeointensity have been reported by Kobayashi (1968) for 2700 Ma samplesfrom the Stillwater Complex, Montana, and by Hale (1987) for 3500 Ma samplesfrom the Komati Formation, Barberton Greenstone Belt, South Africa. However


Figure 24. Summary of available Virtual Dipole Moment (VDM) values from palaeointensitystudies from rocks younger than 360 Ma. VDM values are plotted relative to the present-dayvalue of 8.02× 1022 Am2. CNS, Cretaceous normal superchron; P-CRS, Permo-Carboniferousreversed superchron. Compilation of data from various sources: , Data for the Carboniferous andPermo-Carboniferous from the Liverpool laboratory:4, Data for the CNS from Pick and Tauxe(1993). (after Thomas et al., 1995).

there are no palaeonintensity determinations between 3500–2700 Ma and it is notknown whether the low persisted during that period.

7. Tiny Wiggles

Marine magnetic surveys of fast spreading (>50 mm/yr half rate) oceanic crustoften show a pattern of small scale magnetic anomalies (25–100 nT amplitude,8–25 ka wavelength) superimposed on the more generally recognized spreadingpattern. There has been much discussion on the origin of these “tiny wiggles”.Although some part of them may be due to variations in source layer thickness,geochemistry, or degree of alteration, the linearity and symmetry with respect tothe spreading axis and the consistent pattern of some tiny wiggles on multiplespreading segments, strongly suggests a field related origin. Hartl et al. (1993) havegiven a short account of some of the first investigations. It was generally thoughtthat tiny wiggles were missed, short duration reversals, although the possibility thatthey were the result of palaeointensity fluctuations was suggested.

Cande and Kent (1992b) analysed the magnetic record over the Indian Oceanduring a period of rapid sea-floor spreading in which a clear pattern of tiny wiggles

180 J. A. JACOBS

is observed superimposed on the long reversed intervals between anomalies 24and 27. They also found the same distinctive pattern between anomalies 26 and27 over fast spreading oceanic crust several thousand km away in the southeastPacific. The similarity in detail between these two profiles is strong evidence thatthe tiny wiggles are the result of variations in the palaeomagnetic dipole field. Bycomparing the general characteristics of tiny wiggles with statistical properties ofthe geomagnetic field, Cande and Kent (1992b) concluded that the tiny wiggles arenot likely to be due to complete reversals of the field, but represent a continuousrecord of intensity variations of the palaeodipole field.

D’Argenio et al. (1996) examined in more detail the magnetic anomaly profilesof Cande and Kent (1992a) and, in addition, the raw magnetic anomaly data forthe Cenozoic from similar parts in the south Atlantic. They found that similar anddistinct medium and short wavelength anomalies were present in both the raw dataand in Cande and Kent’s profiles and agreed with Cande and Kent’s conclusionfor their origin. They also carried out spectral analyses which, when convertedinto the time domain using the average spreading rate for each profile, showedthat most of the periodicities appeared to correlate with the Milankovich orbitalparameters expected for that time (Berger et al., 1992) – particularly with the long-term eccentricity periodicities at 3.4, 2.03, 1.2, 0.4 and 0.1 Ma.

Hartl et al. (1993) examined high resolution data from DSD project site 522in the south Atlantic. Smoothing, together with a sampling interval of∼3.5–4.5ka, means that, if short duration features such as the Mono Lake and Laschampexcursions occurred, they would have been missed or largely smoothed beyonddetection. They compared the core locations (in m below sea floor) of the palaeo-magnetic lows from the site 522 record with the temporal locations of C12R tinywiggles deduced by Cande and Kent (1992a,b). Although the correlation is notperfect (missing and/or disturbed record at core breaks, changes in sedimentationrate could account for some of the discrepancy), the number and timing of site 522intensity lows and C12R tiny wiggles are∼ the same.

Gee et al. (1996) compared marine magnetic data for the Central Anomalyobtained over the ultra-fast spreading southern East Pacific Rise with syntheticprofiles based on Brunhes age (0–0.78 Ma) palaeointensities derived from deep-seasediments. The similarity of the synthetic and observed profiles further strengthensthe conclusion of Cande and Kent that the tiny wiggles seen in marine magneticprofiles may largely be the result of changes in the intensity of the geomagneticfield. This interpretation was further supported by systematic variations in thepattern of the Central Anomaly at lower spreading ridges.

The latest support for the interpretation that tiny wiggles in the oceanic mag-netic anomaly record are due to fluctuations in the palaeomagnetic field intensityand not to very short subchrons, has been given by Lanci and Lowrie (1997) whoexamined a 40 m core (the Massicore) drilled through Early Oligocene rocks inthe Marches region of northern Italy. ARM was used as the normalizing parameterfor the NRM. The Massicore contains a complete magneto-stratigraphic record that


matches the geomagnetic polarity sequence from C16n to C12r. The polarity recorddoes not contain any short polarity subchrons that might account for tiny wigglesparticularly within the sequence C13r/C13n/C12r where several are seen in themarine magnetic record. The Massicore record suggests that, if very short sub-chrons exist, their duration is of the order of 10 ka or less, which is very unlikely.Lanci and Lowrie also carried out Fourier analyses of the magnetization profilesof IRM, ARM, NRM and NRM/ARM in the Massicore. The Fourier spectra of theprofiles are noisy, but show significant peaks near the 413 ka eccentricity periodof the Earth’s orbit. Although there is some activity at periods close to the otherMilankovitch periodicities, they are not significant.

8. Summary

Although there has been a significant increase in the last few years in the number ofdeterminations of the palaeointensity of the Earth’s magnetic field, differences be-tween some of the results highlight underlying problems. Absolute intensity valuesobtained from igneous rocks are based on firm theoretical principles and, althoughrigorous experimental procedures are now employed, doubts still remain aboutpossible chemical/mineralogical changes on heating. Sedimentary rocks have theadvantage that they usually contain a continuous temporal record. Relative in-tensities are obtained by normalization of the NRM intensity using one or moreconcentration-dependent parameters. The validity of such normalization dependson the ability to remove the effects of concentration variation without being influ-enced by grain size variation. The possibility that records of sedimentary palaeoin-tensities may be contaminated by climatic variations in material properties is thusstill a matter of concern. If this problem can be resolved, magnetic variations maybe correlated with other climatic parameters and serve as proxy indicators of cli-mate variations. Biological processes and the contribution of biogenic magnetiteto stable remanence also influence magnetic records and a better understanding oftheir effects on palaeointensity determinations needs further investigation.

Apart from these experimental problems, there is the overriding problem thatwe are trying to obtain variations in the intensity of the Earth’s dipole field, andany measurement of the field also contains the non-dipole components which canbe very large. Finally there is the question of chronology – to obtain a true recordof variations of the Earth’s magnetic field, it is essential to have reliable dates forthe records.

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Variations in the Intensity of the Earth's Magnetic Field