Sea-level change and remagnetization of continental shelf ...forth/publications/Oda04.pdf ·...

February 17, 2004 19:1 Geophysical Journal International gji2162

Geophys. J. Int. (2004) 156, 443–458 doi: 10.1111/j.1365-246X.2004.02162.x

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Sea-level change and remagnetization of continental shelf sedimentsoff New Jersey (ODP Leg 174A): magnetite and greigite diagenesis

Hirokuni Oda1,2 and Masayuki Torii31Institute for Marine Resources and Environment, Geological Survey of Japan, AIST, Tsukuba 305-8567, Japan. E-mail: [email protected] Laboratory ‘Fort Hoofddijk’, Faculty of Earth Sciences, Utrecht University, Budapestlaan 17, 3584 CD Utrecht, the Netherlands.E-mail: [email protected] of Biosphere-Geosphere System Science, Okayama University of Science, Okayama 700-0005, Japan

Accepted 2003 October 16. Received 2003 October 13; in original form 2003 February 5

S U M M A R YPalaeomagnetic and rock magnetic studies were performed on a seaward transect of continentalshelf sediments from sites 1071 and 1072 of Ocean Drilling Programme Leg 174A on the NewJersey margin. The sediments recorded a polarity change from reversed to normal, whichwas tentatively interpreted as the Matuyama–Brunhes boundary. The polarity boundary isclosely correlated with the sequence boundary pp3(s), which is considered to have formed asan erosional unconformity during regression and subsequent transgression related to globaleustacy. Rock magnetic analyses indicate that the sediment contains comparable amounts ofmagnetite and greigite (Fe3S4) with possible trace amounts of haematite. A marked change inrock magnetic properties is recorded at Site 1072 with an increase in greigite concentration atthe polarity boundary (62.4 metres below seafloor), where the grain size does not change. AtSite 1071, greigite is dominant from 2 m below pp3(s) to 0.8 m above it. These intervals arecharacterized by higher peak coercivities of remanence and subdued remanence in both themagnetite and greigite components. Palaeomagnetic analysis indicates that the intervals justbelow pp3(s) have dual-component magnetizations of reversed polarity and are considered tohave been deposited during the Matuyama Chron (C1r.1r) and were possibly remagnetizedduring the Brunhes Chron through the formation of the pp3(s) surface. Detrital magnetiteand early diagenetic greigite might have carried the reversed polarity magnetization. Duringthe formation of pp3(s) in the Brunhes Chron, greigite might have formed at the oxidationfront with ongoing downward formation due to oxidation of pyrite by percolating fresh water,which might be a cause of the remagnetization. Our study indicates that careful rock magneticinvestigation is necessary for magnetostratigraphic studies of continental shelf deposits inorder to recognize remagnetizations induced by sea-level changes.

Key words: diagenesis, glacial eustacy, greigite, isothermal remanent magnetization, mag-netic hysteresis, magnetite, remagnetization, sea-level change.

1 I N T RO D U C T I O N

Coring associated with Ocean Drilling Programme (ODP) Leg 174Awas conducted on the New Jersey continental margin to unveil thesedimentary history controlled by eustatic sea-level changes. Strati-graphic interpretation was achieved by means of down-hole loggingand detailed seismic stratigraphy (Austine et al. 1998; Metzger et al.2000; Delius et al. 2001), because sediment core recovery rates werelow due to drilling in shallow water and loss of unconsolidated sandfrom the core barrel.

Sites 1071 and 1072 were drilled on the shelf at depths shal-lower than 100 m. These sites form a seaward transect together withSite 1073 on the edge of the continental shelf (Fig. 1a). The recov-ered sediments consist of clay to sand of Miocene to Holocene age,which were deposited in a shallow marine environment. Several se-

quence stratigraphic boundaries, generated in response to sea-levelfluctuations, were recognized (Austine et al. 1998). These sequenceboundaries are mostly traceable from site to site on the seismicrecords (Fig. 1b). Determining the age of these boundaries was cru-cial for reconstructing the sedimentary history; however, this taskwas not easy due to the paucity and/or reworking of microfossils andto poor sediment recovery. Knowledge of the geomagnetic reversalstratigraphy, coupled with a thorough understanding of the originsof the carriers of the natural remanent magnetization, may help toconstrain the age of deposition.

At sites 1071 and 1072, a Plio-Pleistocene sequence bound-ary pp3(s) separates homogeneous clayey silts below from over-lying thinly interbedded sandy silts and clayey sands (Fig. 2). Thepp3(s) sequence boundary is interpreted as an erosional unconfor-mity formed during regression that was further modified by erosion

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444 H. Oda and M. Torii

Figure 1. (a) Location map of the studied ODP drill sites. (b) Seismic line Oc270 Profile 885 along sites 1071 and 1072. Interpreted sequence boundaries aremarked by thick broken lines.

due to shoreface retreat during subsequent sea-level transgression(Metzger et al. 2000). McCarthy & Gostlin (2000) interpreted pp3(s)to have formed during a low-sea-level stand corresponding to ma-rine oxygen isotope stage 12 (MIS12 ∼ 450–425 ka). This is con-sistent with a large increase in sea level (∼160 m) estimated forthe transition from glacial stage 12 to interglacial stage 11 (Thunellet al. 2002). Benthic foraminifera indicate that the sediments belowpp3(s) were deposited in inner neritic environments (0–50 m) at Site1071, and in upper middle neritic environments (50–65 m) at Site1072 (Austine et al. 1998).

The sequence boundary pp3(s) is almost coincident with areversed to normal polarity boundary and was tentatively in-terpreted as the Matuyama–Brunhes boundary during ODP Leg

174A (Fig. 2). At Site 1072, magnetic susceptibility abruptly de-creases upward at the polarity boundary, which occurs 4.9 m be-low pp3(s). This susceptibility change may be related to a min-eral magnetic change after deposition because grain size does notchange at the boundary, and thus mineral magnetic study is im-portant to constrain the palaeomagnetic polarity at the time ofdeposition.

Some samples from ODP Leg 174A showed evidence for acqui-sition of gyroremanent magnetization (GRM) above 60 mT dur-ing alternating field demagnetization (AFD) (Austine et al. 1998),which might be an indication of the presence of single-domaingrains such as greigite (Snowball 1997a). Greigite has been re-ported in lacustrine sediments (Skinner et al. 1964; Dell 1972;

C© 2004 RAS, GJI, 156, 443–458


Magnetism of ODP site 1071 and 1072 sediments 445

Figure 2. Schematic lithological columns for sites 1071 and 1072. Cored intervals with sediment recovery are shown by black areas. Lithology by ‘inference’is based on cored material and down-hole logging. A Plio-Pleistocene sequence boundary pp3(s) is closely associated with the Matuyama–Brunhes (M − B)boundary, as discussed in this paper. All modified after Austine et al. (1998).

Snowball & Thompson 1990), in rapidly deposited marine sedi-ments that were subsequently uplifted and exposed on land (e.g.Roberts & Turner 1993; Roberts 1995; Horng et al. 1998), andshallow marine Holocene sediments (Lee & Jin 1995). The bacte-rially mediated reduction of sulphate and H2S production in sub-oxic and anoxic sediments control the dissolution of iron oxides(Canfield & Berner 1987). This process is often accompanied by asubsequent conversion of magnetite by way of greigite into param-agnetic pyrite (Karlin & Levi 1983, 1985). Greigite may be pre-

served if the pyritization process is arrested (Roberts & Turner1993).

In magnetostratigraphic studies it is important to understand theeffects of diagenesis on both magnetic iron oxide and iron sulphideminerals in continental shelf sediments whose deposition is con-trolled by sea-level change. The present study was designed to iden-tify the magnetic minerals at sites 1071 and 1072, to investigate thediagenetic effects of sea-level change on magnetic minerals, and toclarify the possibility of remagnetization below pp3(s).

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2 M AT E R I A L S A N D M E T H O D S

Palaeomagnetic samples were taken from clayey to silty sedimentsspanning the tentatively determined polarity boundary that occursnear sequence boundary pp3(s). Discrete samples (7 cm3 cubes)were collected at stratigraphic intervals of approximately 10 cmfrom the working halves of cores from Holes 1071C and 1072A.Anisotropy of magnetic susceptibility (AMS) was measured witha susceptibility anisotropy meter (AGICO Kappabridge KLY-3S) atthe Geological Survey of Japan (GSJ) to check for physical dis-turbance during or after sedimentation or coring. The degree ofanisotropy (PJ ) and the shape parameter (T) were calculated ac-cording to Jelinek (1981).

Subsequently, the natural remanent magnetization (NRM) wasmeasured with a 2G Enterprises superconducting rock magnetome-ter (SRM model 760) at the GSJ. Typical noise levels of the mag-netometer are lower than 1 × 10−12 A m2. All the samples weresubjected to progressive alternating field demagnetization (PAFD)at 8 steps up to 80 mT using a 2G Enterprises demagnetizer thatis configured in line with the SRM. Demagnetization was con-ducted successively along the sample Z-, X - and Y -axes. Aftermeasurement of the NRM, an anhysteretic remanent magnetiza-tion (ARM) was imparted with a DC bias field of 100 µT appliedparallel to the peak alternating field of 80 mT. Isothermal rema-nent magnetizations (IRMs) were imparted with a pulse magne-tizer (2G Enterprises model 660) during stepwise acquisition up to2.5 T. S−0.3T was also calculated according to Bloemendal et al.(1992):

S−0.3T = [−(IRM−0.3T/SIRM) + 1]/2, (1)

where SIRM is the saturation IRM, which was imparted with apulsed magnetic field of 2.5 T, and IRM−0.3T is the magnetiza-tion after the application of a back-field of 300 mT. Kruiver &Passier (2001) pointed out that the new definition of the S ratio byBloemendal et al. (1992) is more logical than the classical definitionof the S ratio by Thompson & Oldfield (1986) because the new defi-nition directly expresses the ratio of saturation remanence carried bylower-coercivity minerals to the total magnetic mineral assemblageprovided that no magnetic interactions occur.

Magnetic hysteresis measurements were made on 12 selectedsamples using an alternating gradient magnetometer (MicroMagmodel 2900, Princeton Measurements Corporation) at the Univer-sity of Utrecht (the Netherlands) with a phenolic probe. Sedimentsamples were weighed, glued in to plastic sample cylinders whilewet and measured immediately, to avoid problems caused by chemi-cal change during drying. Hysteresis loops were obtained by cyclingthe applied magnetic field between +1 T and −1 T. Sample weightswere ∼10–20 mg and the sensitivity of the measurements was∼1 nA m2.

Four samples were selected from each of holes 1071C and 1072Aand were subjected to low-temperature magnetic measurementswith a Quantum Design magnetic property measurement system(MPMS-XL5) at the GSJ. Samples were cooled in zero field to 6 K,at which temperature the 2.5 T DC field was applied for 60 s. Themagnet was then reset to zero. The SIRM was then measured at 2◦

steps during warming up to 300 K.Two samples were selected from each of holes 1071C and

1072A and magnetic minerals were extracted with a magnetic fin-ger (Kirschvink et al. 1992). X-ray diffraction (XRD) analyseswere made on the magnetic extracts using a JDX-8030W instru-ment (JEOL) at the GSJ. Diffractograms were collected between5◦ and 55◦ (2θ ) with Cu Kα radiation at 0.02◦ step scan, with 1 s

measurement time per step. After XRD analysis, thermomagneticanalyses were made with a thermomagnetic balance at Kyoto Uni-versity (noise level of ∼ 2 × 10−7 A m2). The measurements weremade in a DC magnetic field of 0.85 T at a heating/cooling rate of8◦ min−1 in air.

3 R E S U LT S

3.1 NRM demagnetization behaviourand magnetic mineralogy

Due to coring with extended or rotary core barrels, the recoveredsediments underwent relative rotation between adjacent sedimentblocks, and palaeomagnetic declination values are not meaning-ful. Typical examples of PAFD results are shown for samples withstable normal polarity from Hole 1071C in Fig. 3(a) and for sam-ples with stable reversed polarity from Hole 1072A in Fig. 3(b),respectively. Soft magnetic components were generally removedby demagnetization with a peak field of 20 mT. Some samples(Fig. 3c) have a higher-coercivity component with reversed polar-ity and an overlapping normal polarity component. For some othersamples, unstable behaviour is observed (Fig. 3d), where the pri-mary magnetization might have had reversed polarity; however, astable palaeomagnetic direction cannot be determined. Fig. 4(a) isan example of a sample for which a deflection away from the ori-gin is observed above 60 mT in the direction perpendicular to theY -axis (the last axis along which a field was applied during demag-netization treatment at GSJ). Shipboard PAFD treatment revealedthat some samples from holes 1071C and 1072A unambiguouslyacquired a GRM in the direction perpendicular to the Z-axis (thelast axis along which a field was applied during static AFD on theship) above 60 mT and up to 200 mT (e.g. Fig. 4b). GRM hasbeen shown to exist in anisotropic magnetic material composed ofsingle-domain magnetic grains (Stephenson 1980), where the spuri-ous magnetization is produced in the direction perpendicular to theapplied alternating field and anisotropy axes. The general character-istics observed for these samples are similar to the strong GRM ob-served during static three-axis PAFD on greigite (Snowball 1997a;Hu et al. 1998; Sagnotti & Winkler 1999; Stephenson & Snowball2001).

Zero-field warming curves of an SIRM acquired at 5 K for sites1071 and 1072 show the Verwey transition between 107 and 119 K(Fig. 5). The transition temperature is mostly above 110 K, thereforethe composition is close to magnetite without a significant amountof low-temperature oxidation or titanium/aluminium substitution(Kozlowski et al. 1996). Proterozoic low-Ti iron oxide deposits arewidely exposed in New York and New Jersey, including the Adiron-dack Mountains (e.g. Foose & McLelland 1995), from where theHudson River originates. The nearly pure magnetite in the studiedmaterial might have been transported from that region. The pres-ence of detrital magnetite has also been reported for the Pleistocenecontinental slope sediments off New Jersey at ODP sites 903 and904 (Urbat 1996).

The absence of a characteristic magnetic transition at 30–34 K inthe low-temperature SIRM (Fig. 5) suggests that pyrrhotite does notoccur in detectable amounts in these samples (Rochette et al. 1990).XRD results for four selected magnetic extracts from representativeintervals of high SIRM/K ratios for holes 1071C and 1072A containpeaks corresponding to magnetite and greigite (Fig. 6).

Thermomagnetic results indicate the Curie temperature of mag-netite to be at 575–580 ◦C (Fig. 7). A large decrease in magneti-zation is also evident at around 250 ◦C (Fig. 7a). Figs 7(b) and (c)

C© 2004 RAS, GJI, 156, 443–458



1072A-9R2, 6 cm (62.56 mbsf)(b)

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Figure 3. Typical examples of vector endpoint diagrams of progressive alternating field demagnetization experiments. Solid (open) circles denote the projectionof the vectors on to the horizontal (vertical) plane. (a) Normally magnetized sample from Hole 1071C and (b) reversely magnetized sample from Hole 1072Awith a stable magnetization above 10 mT. Samples with unstable features for (c) Hole 1071C and (d) Hole 1072A.

are thermomagnetic curves measured through several minor cyclesto reveal the temperature at which the magnetic minerals change ordecompose. Both curves indicate that below 220 ◦C the thermomag-netic curves are reversible, but between about 220 ◦C and 340 ◦Cthe curves become irreversible. This indicates that a magnetic phasedecomposed between 220 ◦C and 340 ◦C.

It has been reported that greigite decomposes irreversibly above200 ◦C (Krs et al. 1992; Reynolds et al. 1994; Roberts 1995; Toriiet al. 1996). Thermomagnetic curves for lake sediments contain-ing greigite show a decrease of magnetization during heating above300 ◦C, which is similar to that observed in this study (Roberts et al.1996). Dekkers et al. (2000) also reported that synthetic greigitestarts to alter at ∼200 ◦C and that it loses half of its magnetizationbetween 243 ◦C and 276 ◦C during heating in air. This thermomag-netic behaviour combined with low temperature magnetic propertiesand XRD analysis supports the interpretation that greigite is com-mon in the studied sediments in addition to magnetite. Greigite-bearing samples usually have high SIRM/K ratios between 50 and

80 kA m−1 (Snowball 1991, 1997b; Roberts 1995), but greigite-bearing sediments have been found with low SIRM/K values of5 kA m−1. SIRM/K for our samples (see Figs 8 and 9) is above15 kA m−1 for the intervals at 60.5–61.0 metres below seafloor(mbsf) and ∼61.8 mbsf in Hole 1071C, and at ∼62.0 mbsf in Hole1072A. The contribution of paramagnetic minerals to the suscep-tibility might have decreased the ratio (see high-field slope valuesin Table 1). Nevertheless, these intervals with elevated SIRM/Kvalues are considered to represent horizons with higher greigitecontent.

3.2 Hole 1071C

Palaeomagnetic and rock-magnetic results for Hole 1071C are plot-ted versus depth in Fig. 8. The NRM inclinations indicate thatthe polarity boundary occurs at 61.3 mbsf. This polarity bound-ary lies 30 cm below the erosion surface, which is recognized asa sharp lithological boundary between clay (below) and silt/sand

C© 2004 RAS, GJI, 156, 443–458



(a) 1071C-2X2, 125 cm (61.15 mbsf)

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Figure 4. Vector endpoint diagrams of progressive alternating field demagnetization suggesting the presence of a gyroremanent magnetization (GRM) above60 mT. (b) Reproduced from Austine et al. (1998) for sample 1071C-13X1, 59 cm (153.49 mbsf), which shows strong acquisition of a GRM from 60 to 200 mT.

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1071C-2X3, 3 cm (61.43 mbsf)

1072A-9R1, 123 cm (62.23 mbsf)

1072A-9R2, 15 cm (62.65 mbsf)

Figure 5. Results of low-temperature magnetic measurements for holes 1071C, and 1072A. Solid curves (left axis) are the results of zero-field warming of anIRM imparted in a DC field of 2.5 T from 6 K after zero-field cooling. Dotted curves are the derivative of the curves (right axis).

layers (above) at 61.0 mbsf. The NRM intensity increases up-coreabove 61.0 mbsf from about 0.002 A m−1 to around 0.02 A m−1 andincreases again to around 0.07 A m−1 at 60.0 mbsf.

In contrast to the results from Hole 1072A (see Section 3.3below), the magnetic susceptibility does not change across thepolarity boundary in Hole 1071C, although higher suscepti-bilities are observed above 60.0 mbsf. Magnetic anisotropyparameters (not shown here) indicate a primary sedimentary fab-ric (T > 0, K min inc >67◦) throughout the studied interval,which suggests that the sediments have not been disturbed sincedeposition.

Parameters indicative of magnetic mineral concentration (ARMand SIRM) increase above 61.0 mbsf. SIRM values have a peakat 61.8 mbsf, where there was only a small local maximum forARM. S−0.3T is fairly constant at around 0.98 below 61.0 mbsf,and gradually decreases to 0.95 at 58.4 mbsf. This indicates thatthe contribution of higher coercivity minerals, such as haematite, isrelatively small, but that it increases slightly upward. The remarkablylow value of S−0.3T at 60.98 mbsf is 8 cm below the pp3(s) erosionalsurface. SIRM/K values gradually increase up-core, with a peakvalue at 61.8 mbsf, which increases suddenly at 61.0 mbsf to avalue of around 15 × 103 A m−1.

C© 2004 RAS, GJI, 156, 443–458



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Figure 6. X-ray diffractograms for (a) Sample 1071C-2X1, 93 cm and (b) Sample 1072A-9R2, 15 cm. Identified peaks are labelled for magnetite (Mt),maghemite (Mh), greigite (Gr), chlorite (Ch), and quartz (Qz).

3.3 Hole 1072A

Palaeomagnetic and rock-magnetic results for Hole 1072A are plot-ted versus depth in Fig. 9. The NRM intensity suddenly diminishesup-core by one order of magnitude at 62.3 mbsf. The palaeomag-netic inclination is negative below 62.3 mbsf and positive above62.3 mbsf for shipboard measurements on half-core samples. Theinclinations for the discrete samples below 62.4 mbsf are negative;however, the inclinations above 62.3 mbsf are not always positive,but show a general trend toward negative values. The positive in-clination above 62.3 mbsf in the half-core data might result fromincomplete removal of a positive drilling-induced overprint at thelow maximum demagnetization field of 20 mT and from integra-tion of the signal from undisturbed and disturbed outer parts of thehalf core. Discrete samples do not provide clearly defined primarymagnetizations due to the unstable nature of these sediments.

Magnetic susceptibility abruptly decreases up-core above62.3 mbsf. Magnetic anisotropy parameters (not shown here) indi-

cate a primary sedimentary fabric (T > 0, K min inc >67◦) through-out the studied interval, as was also observed for Hole 1071C. Otherconcentration-dependent magnetic parameters (ARM and SIRM)also consistently decrease above 62.3 mbsf. S−0.3T is constant atvalues of around 0.98 below 62.3 mbsf, but it gradually decreases to0.96 at 61.0 mbsf. SIRM/K values are constant at about 5 kA m−1

below 62.3 mbsf, but increase up-core at 62.3 mbsf forming a broadmaximum with a peak value of ∼15 kA m−1 at 62.0 mbsf beforedecreasing to around 6 kA m−1 at 61.1 mbsf.

4 D I S C U S S I O N

4.1 Hysteresis loop analysis

A typical magnetic hysteresis loop for the studied sediments isshown in Fig. 10. For the analysis of hysteresis loops, Jackson et al.(1990) proposed taking the derivative after subtracting the lower part

C© 2004 RAS, GJI, 156, 443–458



of the loop from the upper part with the use of the Fourier transform.Instead of using derivatives, analysis using hyperbolic basis func-tions was proposed by von Dobeneck (1996). This approach usesnot only the remanent hysteretic magnetizations (upper loop minuslower loop divided by two; dotted curve in Fig. 10), but also theinduced hysteretic magnetizations (mean of upper and lower loopsdivided by two; dashed curve in Fig. 10). Hyperbolic basis functionsare more effective than taking the simple derivative when decom-posing the hysteresis curves, especially when overlapping coercivitycomponents are involved. With the use of the analysis program pro-duced by von Dobeneck (1996), basic parameters for the hysteresisloops were also obtained and are listed in Table 1. The program isdesigned to incorporate drift correction, gridding with the use ofsecond-degree polynomials, and centralization (rotate 180◦ along

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Figure 7. Thermomagnetic curves for magnetic extracts. The vertical axesare normalized to the initial value. (a) Sample 1071C-2X3, 3 cm with heatingand cooling cycles, (b) subsample from the same stratigraphic level withminor heating/cooling cycles and (c) Sample 1072A-9R2, 15 cm with minorcycles. The heating curves are reversible up to 220 ◦C, but are irreversiblebetween 220 ◦C and 340 ◦C. Above 340 ◦C, the heating curve shows theCurie temperature of magnetite (∼580 ◦C). During cooling no features wereobserved except the Curie temperature of magnetite.

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Figure 7. (Continued.)

the origin to average) to reduce noise. The most important part ofthe program is the inversion utilizing the hyperbolic basis functionswith a set of characteristic coercive forces. The inversion is designedto find the best combinations of the basis functions with the con-straint of simplest solutions to minimize the number of functions.The spectra of the hyperbolic basis functions for samples from holes1071C and 1072A are shown in Fig. 11. Three fractions are recog-nized with low (phase I; 30–55 mT), medium (phase II; 55–150 mT)and high (phase III; 150–1000 mT) coercivities.

4.2 Decomposition of IRM acquisitioncurves and greigite content

Investigations of the coercivity spectrum were also made by theanalysis of IRM acquisition curves. Robertson & France (1994)showed that the IRM acquired by natural magnetic assemblages canbe expressed by a cumulative log-Gaussian function of the mag-netizing field. McIntosh et al. (1996) used a subjective fit by eyeon the gradient of the IRM acquisition curve versus the log of theapplied field. On the other hand, Stockhausen (1998) introducedgoodness-of-fit parameters for the analyses. These are the residualsum of squares between the model and the data for: (1) the gradientof the IRM acquisition curve, (2) the gradient of the log of the IRMacquisition curve, or (3) the IRM acquisition curve. Kruiver et al.(2001) introduced the standardized acquisition plot as a target resid-ual to be minimized in order to facilitate a better fit at the lower-and higher-coercivity ends in addition to the gradient of log IRMand the IRM acquisition curve. Although both of these methods arerobust, the criteria for which residuals should be prioritized dependon the operator. Heslop et al. (2002) proposed a fully automatedmethod to fit the gradient of the logarithm of the IRM accordingto the maximum likelihood principle (IRMunmix software). Theirmethod is fairly robust and powerful thanks to the EM algorithm(Dempster et al. 1977). An example of the analysis for representa-tive samples from Leg 174A is shown in Fig. 12. The logarithm ofthe IRM acquisition curve (Fig. 12a) and the logarithm of the IRMgradient (Fig. 12b) are plotted along with the residuals of the fits.The optimized solutions were checked by eye for the goodness-of-fit between data and model, both on the plot of log IRM gradient

C© 2004 RAS, GJI, 156, 443–458



Figure 8. Palaeomagnetic and rock magnetic results for Hole 1071C. (a) Palaeomagnetic inclinations for the discrete samples determined by fitting linearregression lines (solid circles) are plotted, together with those of the pass-through measurements (at 20 mT) conducted on the ship (dotted curves). (b) NRMintensity is after demagnetization at 20 mT. Rock magnetic parameters are plotted down-hole for: (c) ARM, (d) SIRM, (e) S−0.3T, (f) volume magneticsusceptibility (K), and (g) SIRM/K, respectively. The right-hand column is a simplified lithological log with the upper legend indicating grain size variations.IW marks the position of an interstitial water sample. PP3(s) is a Plio-Pleistocene sequence stratigraphic boundary (Austine et al. 1998). The horizontal dashedline is the preliminary polarity boundary estimated from the shipboard results.

Figure 9. Palaeomagnetic and rock magnetic results for Hole 1072A. Details are the same as in Fig. 8. The horizontal dashed line is the preliminary polarityboundary determined from the shipboard results.

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Figure 10. A typical example of a magnetic hysteresis loop for sample1071C-2X3, 3 cm (solid curves) after paramagnetic slope correction. Thedotted and dashed curves are the remanent hysteretic magnetization (dif-ference between upper and lower loops) and the induced hysteretic magne-tization (mean of upper and lower loops), respectively, as defined by vonDobeneck (1996).

and IRM acquisition combined with the residuals. In general, twocomponents are enough to produce a reasonable model with smallresiduals. However, their method fails when three-component mod-els were selected for the case of a ‘non-saturated’ solution, as pre-dicted by Heslop et al. (2002). Thus, we use the ‘saturated’ optionfor the analysis described below.

In order to determine the optimum number of parameters, we in-troduce Akaike’s information criterion (AIC) instead of the t-test orF-test, as used by Kruiver et al. (2001) and Heslop et al. (2002),for the optimum number of components. AIC is a measure of dis-tance between the model and observation for any kind of thoughtfulmodel (Akaike 1974). The most likely model is selected as givingminimum AIC. On condition that the standard deviation of the ob-servation error is unknown, AIC is calculated by the approximationwith a large number of data points compared with the number ofparameters (Burnham & Anderson 1998), so that:

AIC = N ln(S/N ) + 2K , (2)

Table 1. Hysteresis parameters for representative samples.

Sample Depth M s M r H c H cr Slope M r/M s H cr/H c

(mbsf) (A m2 kg−1) (A m2 kg−1) (mT) (mT) (A m2 kg−1 T−1)

1071C-2X1, 93–95 cm 59.33 9.09E–02 1.02E–02 13.2 43.0 5.85E–02 0.11 3.261071C–2X2, 67–69 cm 60.57 2.12E–02 3.56E–03 19.0 44.6 5.88E–02 0.17 2.351071C–2X2, 134–136 cm 61.24 1.38E–02 1.74E–03 18.4 57.9 6.35E–02 0.13 3.151071C–2X3, 3–5 cm 61.43 1.84E–02 2.85E–03 17.2 62.9 7.13E–02 0.15 3.661071C–2X3, 74–76 cm 62.14 1.29E–02 1.24E–03 12.1 52.4 5.60E–02 0.10 4.311071C–2X4, 26–28 cm 63.16 1.55E–02 1.22E–03 10.1 45.0 6.27E–02 0.08 4.47

1072A–9R1, 14–16 cm 61.14 1.35E–02 1.07E–03 10.1 51.5 4.38E–02 0.08 5.111072A–9R1, 81–83 cm 61.81 2.08E–02 2.46E–03 14.9 58.4 0.12 3.931072A–9R1, 123–125 cm 62.26 2.14E–02 2.08E–03 12.4 60.8 5.63E–02 0.10 4.891072A–9R1, 133–135 cm 62.33 3.59E–02 3.18E–03 12.5 61.9 4.61E–02 0.09 4.941072A–9R2, 15–17 cm 62.65 1.02E–01 5.27E–03 7.0 40.1 5.86E–02 0.05 5.701072A–9R2, 81–83 cm 63.31 8.67E–02 3.59E–03 5.8 42.1 4.80E–02 0.04 7.24

where N , K and S are the number of data points, number of param-eters and sum of the squared residuals, respectively. When the ratioN/K < 40, the bias adjustment for small sample size is requiredfor the calculation of AIC (Burnham & Anderson 1998), so that:

AICc = N ln(S/N ) + 2K + [2K (K + 1)]/(N − K − 1). (3)

The model is optimal when AICc is minimum among the consid-ered models. AICc was calculated for models with a different num-ber of components, and the results were compared with each other(Table 2). AICc was calculated on each sample using both models,and the models with two peaks are optimal (minimum AICc) in mostcases. Although there are some samples for which the best model is athree-component model, we consistently used two-component mod-els (Table 3). Another reason for choosing two-component models isthat three-component models are unstable with several local minimaof similar likelihood. The peak coercivity of remanence values forthe component with lower coercivity of remanence is between 49 and76 mT, which is considered to be typical of pseudo-single-domainmagnetite. The peak coercivity of remanence values for the compo-nent with higher coercivity of remanence is between 66 and 98 mT.These values are similar to the measured coercivity of remanenceof single domain greigite, which was reported as 60–95 mT for awide range of greigite-bearing sediments (Roberts 1995; Snowball1997a).

From the investigation of magnetic minerals, as presented abovein Section 3, the ratio of high-coercivity magnetizations to the totalmagnetization in the analysis of IRM acquisition curves is consid-ered to represent the relative amount (in terms of magnetization)of greigite in the sample. The results are listed in Table 3 and areplotted versus depth in Fig. 13. Phases I and II in the hysteresisanalyses are also considered to represent magnetite and greigite,respectively. Phase III is assumed to represent haematite, becausedetrital haematite is more likely to occur than goethite in continentalshelf sediments. This is consistent with the fact that IRM acquisitioncurves also show a small portion of unsaturation. The fraction ofphase III is minor, therefore we do not discuss this phase further.

4.3 S−0.07T

Considering the observed coercivity of around 66–98 mT for thehigher-coercivity component, we use S−0.07T as a rough estimate forthe ratio of contribution of the greigite to the total magnetization inthe samples with the following equation:

S−0.07T = [−(IRM−0.07T/SIRM) + 1]/2, (4)

C© 2004 RAS, GJI, 156, 443–458



Figure 11. Plot of results of decomposition into hyperbolic basis functions (von Dobeneck 1996) for (a) Hole 1071C and (b) Hole 1072A. IH and RH representthe induced hysteretic magnetization and remanent hysteretic magnetization, respectively. The histograms denote the percentage contribution of IH or RH to thetotal magnetization of the samples. In general, three magnetic mineral components are recognized. Components I, II and III correspond to magnetite, greigiteand haematite, respectively. See text for further explanation.

where SIRM is the IRM measured after applying a pulsed field of2.5 T along the +Z axis of the sample; IRM−0.07T is measured af-ter imparting a reversed pulsed field of 70 mT along the −Z axis.S−0.07T values are plotted in Fig. 13. The magnetization of the lower-coercivity (magnetite) and the higher coercivity (greigite) compo-nents estimated from IRM decomposition and from S−0.07T are al-most identical. Values of about 0.65 for S−0.07T seem to be a goodthreshold for discriminating between the intervals with higher andlower greigite contents for this study. If S−0.07Tis lower than 0.35,greigite is abundant in the sample. For holes 1071C and 1072A,S−0.07T is lower than 0.65 between 60.1 and 63.1 mbsf, and above62.4 mbsf, respectively. The intervals with higher greigite contentare marked with grey shading in Fig. 13. This interval is mainlycharacterized by subdued magnetizations of magnetite and greigite,and increased coercivity for both components.

4.4 Magnetic mineral diagenesis and time ofmagnetization acquisition

At sites 1071 and 1072, the calcareous nanofossils that mark subzoneCN14a (0.46–0.9 Ma) were found at 64.4 mbsf and 65.0 mbsf, re-

spectively (Fig. 13). At site 1072, calcareous nanofossils that belongto zone CN14 (0.25–0.9 Ma) were found at 57.79 mbsf. The sedi-ments above these horizons with reversed polarity are considered tohave been deposited during the Matuyama Chron (C1r.1r). Althoughit is not clear whether the sediments with doubtful polarity just be-low pp3(s) (Fig. 13; grey horizons) were originally deposited duringthe Matuyama Chron, considering the documentation of negative in-clinations after PAFD at 40 mT (Figs 8 and 9) the sediments mighthave been deposited during the Matuyama Chron (Fig. 14a). Duringearly diagenesis, bacterial sulphate reduction might have facilitateddiagenetic growth of greigite with reversed polarity (Fig. 14b). Also,reductive diagenesis might have resulted in dissolution of magnetite,as was reported by Urbat (1996) for New Jersey continental slopedeposits.

From palynological analysis of sediments from ODP sites 1072and 1073, McCarthy & Gostlin (2000) suggested that unconformitypp3(s) can be correlated from the outer shelf to the upper slope andthat it was generated during a glacio-eustatic sea-level lowstand,probably during marine oxygen isotope Stage 12 (∼450–425 ka).They also suggested that a substantial hiatus is associated with thepp3(s) boundary at Site 1072. McHugh & Olson (2002) provided achronological model for Site 1073 based on oxygen isotope records

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1071C-2X3, 3 cm (61.43 mbsf)

8

6

4

2

0

dIR

M/d

log

(H)

3.02.52.01.51.00.50.0

Log applied field (mT)

0.80.40.0

-0.4

resi

du

al

3.5

3.0

2.5

2.0

1.5

1.0

0.5

0.0

IRM

(A

m-1

)

3.02.52.01.51.00.50.0

Log applied field (mT)

4020

0-20

resi

dual (

x10

-3 A

m-1

)

Data Model

(a)

(b)

Figure 12. An example of fitting with a mixture of two normal distributionson the gradient of the log of the IRM acquisition curve. (a) IRM acquisitioncurve plotted with an abscissa of the log of the applied field. Solid circlesand solid curves represent measured data and the model, respectively. (b)Gradient of the log of the IRM acquisition curve for the measured data (solidcircles) and the model (solid curves). The resolved components are shownas dotted curves with peak coercivities of 76 and 93 mT, respectively.

Table 2. AICc for the IRMunmix analysis with the different number of components.

Sample Depth N SSR AICc

(mbsf) C = 1 C = 2 C = 3 C = 1 C = 2 C = 3K = 3 K = 6 K = 9

1071C-2X1, 93-95 cm 59.33 24 – 1.85E−01 1.91E−01 – −99.79 −85.161071C-2X2, 67–69 cm 60.57 24 3.31E−01 1.45E−01 1.42E−01 −95.64 −105.66 −92.301071C-2X2, 134–136 cm 61.24 24 1.14E+00 1.84E−01 1.91E−01 −65.95 −99.96 −85.171071C-2X3, 3–5 cm 61.43 24 2.30E+00 1.61E−01 1.18E−01 −49.09 −103.24 −96.791071C-2X3, 74–76 cm 62.14 24 1.81E+00 1.71E−01 1.34E−01 −54.82 −101.79 −93.741071C-2X4, 26–28 cm 63.16 24 6.28E−01 1.61E−01 1.65E−01 −80.23 −103.10 −88.63

1072A-9R1, 14–16 cm 61.14 24 1.38E+00 1.99E−01 1.48E−01 −61.33 −98.09 −91.251072A-9R1, 81–83 cm 61.81 24 2.26E+00 2.13E−01 1.01E−01 −49.53 −96.46 −100.371072A-9R1, 123–125 cm 62.26 24 – 2.16E−01 1.18E−01 – −96.09 −96.691072A-9R1, 133–135 cm 62.33 24 2.40E+00 2.38E−01 1.85E−01 −48.08 −93.79 −85.971072A-9R2, 15–17 cm 62.65 22 1.26E+00 1.74E−01 7.73E−02 −55.55 −88.85 −91.331072A-9R2, 81–83 cm 63.31 24 1.17E+00 2.06E−01 2.13E−01 −65.26 −97.27 −82.53

SSR: sum of squared residuals. AICc: Akaike’s information criterion (AIC) with small-number correction (see text for details). N : number of data in IRMgradient curve. C: number of components. K: number of parameters. Bold values for AICc are the minimum AICc values.

and interpreted pp3(s) as a hiatus separating oxygen isotope stages12 and 11. The MIS 12–11 transition represents one of the most se-vere climate changes of the past half-million years (Howard 1997),and sea level rose by about 160 m at the termination of glacialstage 12 (Thunell et al. 2002). Assuming that the sediment belowpp3(s) was deposited during the Matuyama Chron (C1r.1r), a hia-tus of 0.33–0.47 Myr (between 0.9–0.78 Ma and 0.45–0.43 Ma) isexpected at sites 1071 and 1072.

Peak coercivities of remanence, based on decomposition of IRMcurves (Figs 13b and e), for greigite and magnetite are both higherfor horizons with higher greigite content. The coercivities of re-manence are highest near the base of the doubtful polarity intervals(grey horizons). These horizons also have higher contributions fromgreigite in terms of the magnetization. Jiang et al. (2001) suggestedseveral stages for the formation of greigite by progressive diagenesisafter deposition. They inferred that oxidation of pyrite released fer-ric iron, which was used for the late diagenetic formation of greigitein the vicinity of pyrite. Weaver et al. (2002) provided evidence thatmagnetic iron sulphide minerals, in their case pyrrhotite, can growduring late sedimentary diagenetic reactions as a result of fluid flowevents. On the New Jersey continental shelf, horizons with highercoercivities for greigite and magnetite could represent a trapped ox-idation front (Fig. 14d) produced by downward percolation of oxicfresh water from the overlying porous sediments. This chemical frontmay have facilitated maximum production of secondary greigite byoxidation of pyrite during formation of pp3(s). Sediment grain sizeis coarser at Site 1072 (silt) than at Site 1071 (clay). The greaterporosity of the coarser-grained sediments at Site 1072 would haveenabled more rapid and deeper penetration of oxic fresh water dur-ing formation of pp3(s). In the sediments above the oxidation frontand below pp3(s), especially at Site 1072, the secondary greigitethat formed during progression of the oxidation front might havebeen partially dissolved, together with early diagenetic greigite, as aresult of more oxic conditions which are not favourable for greigitepreservation.

At Site 1072, the lower amount of magnetization carried by mag-netite for the upper horizon may indicate that progressive oxidationof magnetite occurred, followed by suboxic conditions around theoxidation front that may have caused the effective removal of ferriciron by oxic fresh water during formation of pp3(s). It is also possi-ble that highly permeable sand layers allow present-day fresh waterto percolate via ground water connections to the continental shelf

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Table 3. Parameters of peaks for the log-normal distributions calculated using IRMunmix (Heslop et al. 2002)

Two-peak model

1st peak 2nd peakDepth SIRM High-coercivity

Sample (mbsf) (A m−1) IRM (A m−1) B1/2 DP B1/2 (mT) IRM (A m−1) B1/2 DP B1/2(mT) ratio in M

1071C-2X1, 93–95 cm 59.33 11.21 7.73 1.75 0.33 55.7 3.48 1.87 0.16 74.0 0.311071C-2X2, 67–69 cm 60.57 5.96 3.92 1.74 0.31 55.6 2.04 1.82 0.17 65.7 0.341071C-2X2, 134–136 cm 61.24 2.50 1.49 1.81 0.33 65.0 1.01 1.92 0.14 83.2 0.401071C-2X3, 3–5 cm 61.43 3.70 1.73 1.88 0.32 76.2 1.97 1.97 0.12 93.3 0.531071C-2X3, 74–76 cm 62.14 1.89 1.22 1.80 0.32 63.3 0.67 1.96 0.11 91.7 0.351071C-2X4, 26–28 cm 63.16 1.84 1.27 1.70 0.34 50.7 0.57 1.88 0.16 76.6 0.31

1072A-9R1, 14–16 cm 61.14 1.77 1.22 1.77 0.34 58.7 0.55 1.97 0.12 92.5 0.311072A-9R1, 81–83 cm 61.81 3.41 1.84 1.85 0.35 70.7 1.58 1.99 0.13 98.0 0.461072A-9R1, 123–125 cm 62.26 2.56 1.34 1.81 0.35 64.5 1.22 1.99 0.13 96.7 0.481072A-9R1, 133–135 cm 62.33 8.10 3.65 1.86 0.34 72.8 4.45 1.97 0.13 92.3 0.551072A-9R2, 15–17 cm 62.65 5.73 3.74 1.69 0.33 49.3 1.99 1.93 0.13 84.2 0.351072A-9R2, 81–83 cm 63.31 5.26 3.74 1.73 0.35 53.4 1.52 1.93 0.13 84.4 0.29

B1/2: peak coercivity. DP: dispersion parameter (both are in log10 scale).

Figure 13. Down-hole plots of the results of IRM decomposition and S−0.07T for holes 1071C and 1072A. Water depths of the sites are shown in brackets.(a, d) High-coercivity ratio and S−0.07T versus depth. The high-coercivity ratio (top axis; open triangles) is calculated by fitting model functions on the IRMacquisition curve (Table 2), and S−0.07T (bottom axis; solid circles) is given by eq. (4). Note the inverted axis for S−0.07T, because S−0.07T is unity when thehigher-coercivity content is zero. The threshold of 0.65 is shown as vertical dashed lines for S−0.07T. (b, e) Solid (open) triangles represent coercivities ofremanence calculated by model fitting on IRM acquisition curves for the high (low) coercivity component. (c, f) IRM intensity for the high (low) coercivitycomponent calculated by model fitting on IRM acquisition curves is plotted as solid (open) triangles. The intensity of the IRM for the high (low) coercivitycomponent calculated from S−0.07T are also plotted as solid (open) circles. Samples from the calcareous nanofossil subzones CN14 (0.25–0.9 Ma) and CN14a(0.46–0.9 Ma) are also shown with arrows in the left-most columns. Palaeomagnetic polarities are shown in vertical columns (black = normal polarity, white =reversed polarity). Grey shading indicates the horizons identified as having normal polarity during Leg 174A, but which are suspected to have been depositedduring the Matuyama Chron. Diagonally striped horizons have no data. Shaded intervals correspond to zones with higher greigite contents. The horizontalthick dashed lines represent the unconformity pp3(s).

C© 2004 RAS, GJI, 156, 443–458



Mt

deposition

(post-) depositionalremanence (R)

Matuyama (C1r.1r)(a)

early diagenesis

Gr

authigenicgrowth (R)

Py

GrPy

dissolution

Matuyama (C1r.1r)(b)

low-T oxidation

oxidationfront

fresh water

GrFe3+

Py Gr

erosion

Py Gr

later stagegreigite (N)

low stand

Brunhes(c)

Present day

oxidationfront

Gr

Fe3+

Gr

pp3(s)

Py Gr Gr

dissolution

(d)

Mt Mt

Mt Mt

Mt

Mt

Mt

Figure 14. Cartoon showing the possible mechanism of diagenesis and remagnetization for the studied sediments. Mt, Gr and Py are magnetite, greigite andpyrite, respectively. (a) Acquisition of (post-) depositional remanence carried by magnetite during the Matuyama Chron (C1r.1r). (b) Early diagenesis duringthe Matuyama Chron (C1r.1r) causing authigenic growth of greigite with reversed polarity. Authigenic formation of pyrite and dissolution of magnetite arealso depicted. (c) Downward progression of an oxidation front by percolation of oxic fresh water during a major sea-level low stand in the Brunhes Chron. Atthe oxidation front, later stage authigenic greigite is inferred to have formed in the vicinity of pyrite with normal polarity magnetization. A maghemite skinformed on magnetite grains as a result of the low-T oxidation. (d) As the oxidation front progressed downward, early and later diagenetic greigite was partiallydissolved. Surface maghemite skins on magnetite grains were also dissolved.

(e.g. Groen et al. 2000). The higher coercivity for the magnetitein and around the zone of uncertain polarity can be interpreted asresulting from enhanced stress at the surface of magnetite grainsdue to low-temperature oxidation (e.g. Cui et al. 1994; van Velzen& Zijderveld 1995).

Snowball & Thompson (1990) and Sandgren & Snowball (2001)found high concentrations of greigite associated with marine incur-sions into lacustrine settings. In both cases, the greigite probablyformed due to the presence of sulphide-rich porewaters in brack-ish conditions. While these settings are different from those of thepresent study, the mechanism for greigite formation could be sim-ilar in that restricted sulphidic porewaters were present for limitedamounts of time, which could have contributed to arrested pyriti-zation and preservation of greigite. In British Early Pleistocene es-tuarine clays (Hallam & Maher 1994) the magnetization carried bygreigite has reversed polarity in the middle of the clay unit, whereasthe coarser-grained upper and lower margins of the unit have normalpolarity. Hallam & Maher (1994) interpreted the reversed polaritycarried by greigite to be an early diagenetic (syndepositional) sig-nal, which was later altered and overprinted due to oxidation duringthe Brunhes Chron. The remagnetized normal-polarity sedimentsare characterized by weak magnetizations and low susceptibilities,which is similar to the zone of higher greigite content with weakermagnetization at Site 1072. Although the present-day situations aredifferent, the British estuarine clays can be considered as a similarsetting to the New Jersey continental shelf sediments, especially atSite 1072, which was nearly at the shoreface during formation ofpp3(s).

5 C O N C L U S I O N S

Palaeomagnetic and rock magnetic studies were performed on con-tinental shelf sediments from ODP sites 1071 and 1072. These sed-iments record a reversed to normal polarity transition, which wastentatively interpreted as the Matuyama–Brunhes boundary. Thepolarity boundary is closely correlated with the sequence boundarypp3(s), which is considered to have formed as an erosional unconfor-

mity during glacio-eustatic regression and subsequent transgressionof sea level. Concerning the origin of magnetic minerals and remag-netization mechanisms, the following issues have been clarified byour study.

(1) Primary magnetite and secondary greigite are the dominantremanence carrying minerals.

(2) The sediment in the remagnetized zone with uncertain polar-ity below pp3(s) might have been deposited during the MatuyamaChron (C1r.1r). Any detrital magnetite would have acquired a (post-)depositional remanent magnetization with reversed polarity. Earlydiagenetic greigite might have been created as a result of sulphatereduction and a chemical remanent magnetization with reversed po-larity could have been acquired. Subsequently, during formation ofpp3(s), magnetite and greigite might have been dissolved belowpp3(s) due to oxidation and effective removal of ferric iron by freshwater that penetrated downward from pp3(s). Additional greigitemight have formed below pp3(s), especially at the base of the zonewith uncertain polarity, which might have resulted from greigite for-mation via oxidation of pyrite. At the base of the zone with uncertainpolarity, the coercivity of magnetite probably increased as a result ofincreased stress via cracking of the surface of the magnetite grains.

(3) The parameter S−0.07T (see eq. 4) is a useful indicator of greig-ite concentration. S−0.07T values lower than 0.65, or magnetizationshigher than 0.35 determined by the IRMunmix software, may beused as a threshold for a higher content of greigite.

(4) Magnetostratigraphic studies of continental shelf depositsneed to be conducted with special care because glacio-eustaticallycontrolled remagnetization might affect the palaeomagnetic recordof such sediments.

A C K N O W L E D G M E N T S

Samples were provided through the Ocean Drilling Programme.The authors are indebted to T. Yamazaki and M. Dekkers for com-ments, and A. Roberts and I. Snowball for critical reviews of themanuscript. The authors are grateful to N. Ishikawa for the use of histhermomagnetic balance and to T. von Dobeneck for providing the

C© 2004 RAS, GJI, 156, 443–458



hysteresis analysis program Hystear. The authors thank A. Kosterovfor suggestions on experiments. D. Heslop let us use the IRMun-mix program and gave valuable comments on statistical analysis.HO is partly supported by the Japan Society for the Promotion ofScience.

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