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Continuous record of geomagnetic field intensity between 4.7 and 2.7 Ma from downhole measurements
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Earth and Planetary Science Letters 136 (1995) 541-550
Continuous record of geomagnetic field intensity between 4.7 and 2.7 Ma from downhole measurements J. Thibal a, J.-P. Pozzi a, V. Barth&s b, G. Dubuisson a a EC& Normale Sup&rieure. Luboraroire de Giologie, URA 1316 CNRS, 24 rue Lhomond, 75231 Paris Cedex 05, France h Centre Etudes Nu&aires, Laboratoire d’Electronique de Technologie et Instrumentation, 38041 Grenoble, France
Received 30 May 1995; revised 30 August 1995; accepted 1 September 1995
Abstract A continuous record of the geomagnetic field intensity from 4.7 to 2.7 Ma has been obtained, together with a precise magnetostratigraphy, from downhole magnetic measurements at Site 884 of ODP Leg 145 in the North Pacific. The record confirms the saw-tooth pattern of geomagnetic field intensity proposed by Valet et Meynadier [lo]. Reversals are characterized by a steep intensity decrease followed by a quick regeneration. Over each polarity interval, rapid variations are superimposed over a progressive decrease in the mean intensity. We find that the duration of each polarity interval is inversely proportional to the mean rate of decrease in the field intensity over this period, and thus this duration seems to be pre-determined.
1. Introduction Five years ago, LET1 (French Atomic Agency, CEA), CFP TOTAL and ENS-CNRS initiated a programme with the help of the Schlumberger Company to identify magnetostratigraphy in sediments by using in-situ downhole measurements. The feasability of in-situ magnetostratigraphy was first demonstrated within the weakly magnetized sediments of the Paris Basin [ll. We used a total magnetic field tool with a nuclear resonance magnetometer. To remove the effects of geomagnetic variations, a reference magnetometer was set up far from the disturbance of the rig (at sea, the nearest permanent landbased magnetic observatory is used). The induction measured within the borehole results from both induced and remanent magnetization of the rocks. A
susceptibility log obtained with a dipole-dipole susceptibility tool is then recorded (both tools are now combined). The processing of the combined induction and susceptibility data gives the polarity of the remanence. The induction B(z) generated in the borehole (as a function of depth) by magnetization J(z) is given by Bii( z) = Cij . Ji( z>, where C is a secondorder symmetrical tensor [Z-4] and J(z) is the total magnetization expressed by J(z) = Ji( z) + J,(z) (where Ji and J, are the induced and remanent magnetization). To a first order of approximation, the nuclear resonance magnetometer continuously measures the projection B’(z) of B(z) in the direction of the present-day geomagnetic field. Part B:(z) of B’(z), which is due to the remanent magnetization J,(z), reverses during reversals while Bi, which
0012-821X/9.5/$09.50 0 1995 Elsevier Science B.V. All rights reserved SSDI 0012-821X(95)00173-5
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2. Downhole magnetostratigraphy
is due to the induced magnetization Ji( z> = x( z> . F, does not (x(z) is the susceptibility and F is the intensity of the present-day geomagnetic field). Bi( z) is calculated from the susceptibility log and is then subtracted from B’(z) to obtain B:(z) and the polarity of J,(z). J,(z), the natural remanent magnetization, can include viscous magnetization and remagnetizations. In these conditions and for a geological interval without large directional changes in J,(z), B:(z) is proportional to the intensity of J,(z). During ODP Leg 145, in the Aleutian Sea, magnetic logging was run at several sites. The scientific goal of the leg was a study of the paleoceanography and paleoclimatology of the North Pacific Ocean for the past 10 m.y. Fig. 1 shows the location of the sites drilled, and especially the deepest of the three sites (884) situated on the northern extremity of the Emperor Seamounts, on the Meiji Tongue [5]. In this paper we will show that the geomagnetic field paleointensity can be obtained from the magnetic logging. We present the geomagnetic field intensity variations together with the precise magnetostratigraphy obtained in the same hole (884E), situated at 51”27.034’N, 168”20.216’E. The samples recovered within a hole drilled a few metres away have been used to control the homogeneity of the magnetic mineralogy.
The magnetostratigraphic interpretation at hole 884E 161 is shown in Fig. 2. The susceptibility log (Fig. 2a) and the magnetometer record corrected for the pipe effect (Fig. 2b) are interpreted in terms of a polarity sequence (Fig. 2~). A comparison between this sequence and the geomagnetic polarity timescale from Cande and Kent [7] shows a perfect agreement down to 450 m below the seafloor (bsf), which corresponds to an age of more than 8 Ma. Consequently, the drilled formations down to 450 m bsf and the magnetic logs are dated. There is good agreement between the interpretation of logs for Hole 884E (Fig. 2c) and the polarities deduced from the core mesurements for hole 884B (Fig. 3).
3. Magnetization
and geomagnetic
paleointensity
Paleointensity studies rely on the assumption that the natural remanent magnetization (NRM) of sediments is mainly due to particle alignment under the effect of the torque exerted by the paleogeomagnetic field Fp during the sedimentation process. The intensity of the natural remanent magnetization J, can be expressed by: J, (A/m)
(
= A . C, . Fp (A/m)
(1)
1
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Fig. 1. Location map of the North Pacific Ocean showing Leg 145 drill sites.
1200
J. Thibal et al. / Earth and Planetary Science Letters 136 (I 995) 541-550
B’(z) NT) -0
Geomagnetic polarity time scale
Downhole polarity sequence
where A is a dimensionless coefficient which depends on the nature of the magnetic minerals, and on the size and shape of the grains of these minerals. C, is the concentration of ferromagnetic minerals. Provided A is constant over the considered logged section, one can normalize the concentration C, of
e 0
543
Polarity Inclination
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Fig. 2. Magnetostratigraphic results at Hole 884E. (a) Susceptibility log from dipole-dipole probe below the casing. (b) Borehole induction from the nuclear resonance magnetometer due to both remanent and induced magnetization of sediments (nT). (c) Interpretation as a downhole polarity sequence. Black/white = normal/reverse polarities. (d) Comparison with the geomagnetic polarity timescale of Cande and Kent [7].
350
400
Fig. 3. Comparison between the polarity sequence deduced from the inclination of the remanent magnetization measured on cores at Hole 884B and the downhole polarity sequence.
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magnetic minerals. We limited our first study to the interval 4.7-2.7 m.y. (259-128 m bsf). After 2.7 m.y., the magnetic susceptibility displays large variations (Fig. 21, due to ash and dropstone layers becoming frequent in the sedimentation as a consequence of large climatic changes [5].
4. Rock magnetism experiments The following rock magnetism experiments on 70 samples demonstrate that the magnetic mineralogy is homogeneous over the selected section and that the concentration of magnetic minerals is well repre-
Science Letters 136 (1995) 541-550
sented by the variations in the low field susceptibility. 4.1. Magnetic minerals and grain size Cycles of hysteresis have been conducted on seventeen samples up to 6 T and at two temperatures (- 83°C and 20°C respectively), below and above Morin’s transition characteristics of haematite. The results are identical, demonstrating that haematite does not exist in detectable amounts and cannot be detected by Morin’s transition in this type of oceanic sediment. A typical result is shown in Fig, 4 (sample 164). Fig. 4a shows the - 0.1 to 1 T part of the cycle
Sample 164
Sample 225
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Fig. 4. Two typical examples of high field measurements. Sample 164 is typical of the 131 m of the section, except for the interval 180.5-196 m bsf. Sample 225 is typical of interval 180.5-196 m bsf, showing a large proportion of a haematite-like mineral, and is not considered further.
J. Thibal e? al. / Earth and Planetary Science Letters 136 (1995) 541-550
0.5
Hole
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This diagram to obtain an estimate The results are given grain size is quite single-domain (PSD) 4.2. Concentration
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Fig. 5. Grain-size diagram deduced from hysteresis properties. J,, /J, vs. H,, /H, [9]. The fourteen samples taken over the considered section are confined in the pseudo-single-domain (PSD) range.
which has been conducted up to 6 T. Fig. 4b illustrates the ferromagnetic part of the cycle when the paramagnetic contribution is subtracted. The saturation magnetization J, is obtained for 1.8 T and the coercive force H, is found to be 11.4 mT. Fig. 4c gives the saturation isothermal remanence magnetization (IRM) 1, and the coercive force of the remanence H,. The saturation of IRM is reached at 0.3 T. For all specimens (fourteen) represented by this typical curve, H, ranges from 10.9 to 26.2 mT with a mean value of 14.6 mT and H, ranges between 31.7 and 80.2 mT with a mean value of 43.1 mT. These measurements show that fine-grained magnetite is bound to be the main magnetic mineral. Nevertheless, in the defined depth interval 196180.5 m bsf, the susceptibility log (Fig. 2a) displays very low values, in most cases lower than the minimum value for the probe pre-set for the particular logging operation in question. Fig. 4 shows the high field measurements of sample 225, which is representative of the specimens sampled in this zone. Fig. 4d-f refers to the same measurements as Fig. 4a-c. The results on samples from this interval show the occurrence of a high-coercivity mineral, probably haematite. The fact that no difference has been detected in the hysteresis parameters obtained at the two temperatures above and below Morin’s transition may be due to a content of very fine grained haematite [8]. The grain size of the magnetic mineral is determined from the classical diagram for J,,/J, vs.
545
is well adapted in this case [9] of the grain-size homogeneity. in Fig. 5, and indicate that the homogeneous in the pseudorange.
Cr and susceptibility
Anhysteretic remanent magnetization (ARM) is currently used in paleointensity studies to normalize J, to the variation in the ferromagnetic mineral concentration Cr. We first compare the alternating field demagnetization curves for NRM and ARM. Typical results are given in Fig. 6 and show that the two demagnetization curves are similar. Then we compare the ARM and low field susceptibility x. The crossplot of the two parameters (Fig. 7) indicates that the relationship between susceptibility and ARM is close to proportionality over the section. Therefore, we can use the downhole susceptibility x to normalize the variations in Cr. The above considerations lead us to propose that the paleointensity variations in the geomagnetic field can be obtained from downhole magnetic measurements: from the total field log B’(z) we subtract the induced part B:(z), and so we obtain a signal for B$ z) proportional to J,(z). The effect of the variation in the ferromagnetic mineral concentration C, is normalized to a constant concentration using B[( z).
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Field (mT) Fig. 6. Typical plot showing alternating field demagnetization of both natural remanent magnetization (NRM) and anhysteretic remanent magnetization (ARM; 100 mT, 0.1 mT). The similarity of the two curves shows that the same magnetic grains are involved in the two processes.
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and Planetary Science Letters 136 (1995) 541-550
Thus, the variations in the geomagnetic field paleointensity can be expressed as [B’(z) - BI(z)]/B:(z), which can be obtained from the magnetic logs.
5. Paleointensity record
03 0
100
200
300
400
500
600
700
600
Susceptibility (~10~ SI) Fig. 7. ARM after a 20 mT alternating field demagnetization vs. low field susceptibility x for single samples. The linear relationship between the two parameters indicates that the low field susceptibility is a suitable normalizing parameter for the NRM. It also confirms a uniform magnetic granulomety over the considered section.
2.7
3.3
3.5
The resulting relative paleointensity record is given in Fig. 8, with positive values for normal polarity (black) and negative values for reverse polarity (white). The sampling depth interval of the magnetic tool is 0.15 m, which corresponds to a maximum of 3.5 k.y. and a minimum of 1.5 k.y. respectively for the lowest and the highest values of the accumulation rate deduced from the magnetostratigraphy over the selected section (43 m 1 m.y.-’ and 96 m . m.y. - ‘1. The NRM/ARM ratios measured on specimens after alternating field demagnetization up to 20 mT are shown in the same figure for comparison. There is a good agreement
3.7
3.9
4.1
4.3
4.5
4.7
Age (MY) Fig. 8. Relative variation of the geomagnetic field intensity over the 4.7-2.7 Ma interval obtained from downhole magnetic records at ODP Site 884. The graph is shown in a polarized form (positive and negative values for normal (black) and reverse (white) polarity periods). 0 = NRM/ARM ratios mesured on specimens after AF demagnetization up to 20 mT.
J. Thibal et al./Earth
and Planetary Science Letters 136 (1995) 541-550
4.7-2.7 Ma, confirm the saw-tooth pattern presented by Valet and Meynadier [lo] and extend the record of continuous paleointensity 0.7 m.y. beyond their record. In addition, Valet and Meynadier’s record is representative of field behaviour at low, equatorial latitudes, whereas our result is the first obtained at high latitude. Note that from Valet and Meynadier’s data, the mean virtual axial dipole moment is 4.1 f 2.0 X lo** A m* for normal periods and 3.3 f 1.5 X lo’* A m* for reverse periods. The mean normal and reverse paleointensities calculated for 4.7-2.7 Ma at Site 884 are equal to within a few percent. In both our and Valet and Meynadier’s study the same paleointensity method is applied (although on cores by Valet and Meynadier and on the surrounding sediments within the borehole in our case). Cores can be remagnetized in the core barrel assembly [ 131. A further remagnetization is created during storage of the core while minerals reach their new physical and chemical equilibrium. In addition, coring certainly remagnetizes the wall of the borehole by the effect of the field produced by the drilling and coring device, probably enhanced by stress relaxation and piezomagnetic effects [14]. However, the remagnetized volume of the surrounding rock is very small compared to the tool investigation volume, and the remagnetized volume is, therefore, relatively much larger in the core. Yet, the comparison between the two records shows that re-
between laboratory and downhole measurements. This comparison validates our method based on rapid data acquisition. It has the advantage of not being dependent on core recovery and does not require demagnetization. The paleointensity curve is interrupted over the two ash-dominated intervals from 4.12 to 4.10 m.y. and from 3.50 to 3.49 m.y. Between 3.75 and 3.56 m.y. (respectively 196 and 180.5 m bsf) a highcoercivity haematite-like mineral is dominant, whereas elsewhere magnetite is dominant. We therefore suspect that the coefficient A (Eq. 1) changes. We have no explanation for this localized change in the proportion of magnetic minerals, except that it corresponds to a sharp decrease in clay content [5] which could have enhanced the permeability and favoured fluid circulation. Several characteristic features of the geomagnetic field can be shown on the curve. During the interval of constant polarity, the behaviour of the field intensity is characterized by large variations superimposed on a steady decrease. Each reversal seems to be the result of the coupling of the rapid variations with the progressive decrease in the mean intensity and is followed by a rapid regeneration of the field strength. These features confer on the curve an asymmetrical saw-tooth shape. The only available previous study [lo-121 has shown a synthetic record obtained from several equatorial Pacific and Indian Ocean cores. Our results, which span the period
2.7
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3
3.1
3.2
54-l
3.3
3.4
3.5
3.6
3.7
3.8
3.9
4
Age(My)
Fig. 9. Comparison between Valet and Meynadier’s relative paleointensity curve established from core measurements from the Indian and Pacific Oceans at low latitude and our curve (heavy line) deduced from downhole magnetic records at ODP Site 884 at high latitude for the interval 4-2.7 Ma. Curves are plotted with a common age scale and a positive constant is added to our paleointensity.
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J. Thibal et al./ Earth and Planetary Science Letters 136 (1995) 541-550
magnetization probably does not affect their quality because the cores have been demagnetized in the laboratory. Fig. 9 shows the detailed comparison of the two records along the common period 4.7-2.7 Ma. To make the comparison with Valet and Meynadier’s results easier, we have tuned our record to the same age scale [15] and a positive constant has been added to the relative paleointensity. The similarity in the records is excellent (even in detail), and very satisfying correlations can be established. Note, however, that slight discrepancies in age within constant polarity periods cannot be avoided. These result from specific time distortion of each record due to rapid variations in accumulation rate that cannot be uncovered by magnetostratigraphy. In addition, let us also consider the effect of a lock-in depth of natural remanence in the paleofield-remanence relationship. Such an aquisition depth of post-depositional remanent magnetization constitutes a zone that can be seen as the result of a mixing of normal and reverse grains just below a reversal [ 16-181, but also as a mixing of grains magnetized to varying intensity throughout the polarity periods between reversals. This it affects the signature of intensity variations during periods of contant polarity. This lock-in zone, which is assumed to have a constant thickness, represents a time dura-
2.7
2.9
3.1
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3.5
tion which depends on the accumulation rate. The zone acts on the sedimentary record of paleointensity as a high-frequency filter whose cut-off frequency is proportional to the accumulation rate. The effect of this zone must not be underestimated. For the lowest accumulation rate at Site 884, a lock-in zone of 0.15-0.30 m in thickness as currently estimated in deep-sea sediments [19-211 has a duration of 7 k.y. A detailed study of the effect of the lock-in zone, which would require comparison with other paleointensity records in holes with different sedimentation rates, is unfortunately beyond the scope of this paper. We next study polarity periods excluding transitions themselves, and analyse the decrease in mean field intensity during each period, without attempting to describe in detail the rapid and large variations superimposed over this decrease. Fig. 10 shows the same data as Fig. 8 in an unpolarized form (i.e., not taking polarity sign into account). Valet and Meynadier [lo] suggested that the polarity chron length is related to the regeneration of the field intensity following a transition. In other words, polarity length is proportional to the field intensity jump across the transition. However, some bias can arise from the difference between the mean dipole field intensity of the normal and reverse polarities.
3.7
3.9
4.1
4.3
4.5
4.r
Age (MY) Fig. 10. The relative variations form. A linear fit is calculated
of the geomagnetic field intensity over the 4.7-2.7 Ma interval at ODP Site 884 are plotted in an unpolarized within each polarity Period between 4.7 and 2.7 Ma after removing transitions.
J. Thibal et al. / Earth and Planetary Science Let&n
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I
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0.3
0.4
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Duration of polarity intervals P (MY) Fig. 11. Inverse linear relationship between the rate of decrease of the mean relative paleointensity over a polarity period and the duration of the corresponding period. R is expressed in percent per 0.1 m.y. of the average intensity calculated over the 4.7-2.7 Ma interval. The error bars on the duration of polarity intervals, ranging between 7 and 13%, are deduced from errors in the timescale assumed by Cande and Kent from the duration of chrons [7]. The R error bars are proportional to the residuals of the linear approximation.
Our record leads us to propose another relationship. We have calculated the mean rate R of decrease in the intensity over each polarity interval, using a linear curve fitting (the only one justified by the data), after removing the transitional data within a 40 k.y. window centred qn the reversal (Fig. 10). It seems that the length of a polarity period P is related to the linear rate of decrease R in the field intensity. The relationship between P and R is shown in Fig. 11, which shows that P is inversely proportional to R.
6. Conclusion This case study confirms the possibility of obtaining a precise magnetostratigraphy from continuous downhole magnetic measurements [l]. Our main new result is that a continuous paleointensity record of the Earth’s magnetic field can be obtained from a single logging operation in a deep-sea sedimentary sequence. This was successfully done at ODP Site 884 in the North Pacific. This new record of paleointensity, obtained for a high latitude, is in excellent
I36 {I 995) 541-550
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agreement with the record of Meynadier et al. [l 11 deduced from core mesurements at low latitudes. Each reversal seems to be the result of the coupling of the rapid variations with the progressive decrease of the mean intensity. Reversals are characterized by an intensity decrease followed by a rapid regeneration of the field strength. The saw-tooth pattern is confirmed and extended by 0.7 m.y. During periods of constant polarity, the paleointensity shows rapid variations superimposed on a slow decrease in the mean field. We also propose a new law for the field intensity behaviour between reversals: It seems that the length of a polarity interval is inversely proportional to the mean rate of decrease in the field intensity during this interval. This leads to the surprising result that the duration of a given polarity interval could be immediately determined after the previous reversal.
Acknowledgements We thank X. Le Pichon and J.-P. Valet for helpful comments and J. Pocachard and A. Etchecopar for their contribution to this programme. The high field measurements were conducted with the help of N. Bontemps and P. Monod in the Physics Department of the ENS using a Squid magnetometer. The data used for this study are available on request from the BRG of the LDEO. This work was partly supported by Schlumberger-EPS and by CNRS-INSU (Programme Giosciences Marines). [ PT]
References 111J.-P. Pozzi, V. Barthes, J. Thibal, J. Pocachard,
M. Lim, T. Thomas and G. Pages, Downhole magnetostratigraphy in sediments: comparison with the paleomagnetism of a core, J. Geophys. Res. 98, 7939-7957, 1993. 121J.-P. Pozzi, J.-P. Martin, J. Pocachard, H. Feinberg and A. Galdeano, In-situ magnetostratigraphy: interpretation of magnetic logging in sediments, EarthPlanet. Sci. Len. 88, 357373, 1989. in [31 Y. Gallet and V. Courtillot, Modeling magnetosuatigraphy a borehole, Geophysics 54, 973-983, 1989. [41 Y. Hamano and H. Kinoshita, Magnetization of the oceanic crust inferred from magnetic logging in Hole 395A, Proc. ODP, Sci. Results 106-109, 1990. 151 D.K. Rea, LA. Basov et al., Proc. ODP, Init. Rep. 145, 1992.
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[6] G. Dubuisson, J. Thibal, V. Barthes, J. Pocachard and J.-P. Pozzi, Downhole magnetic logging in sediments during Leg 145: usefulness and magnetostratigraphic interpretation of the logs at site 884, Proc. ODP, Sci. Results 145, in press. [7] S.C. Cande and D.V. Kent, A new geomagnetic polarity time scale for the late Cretaceous and Cenozoic, J. Geophys. Res. 97, 13917-13951, 1992. [8] F.D. Stacey and SK. Banerjee, The Physical Principles of Rock Magnetism 5, Elsevier, New York, 1974. [9] R. Day, M. Fuller and V.A. Schmidt, Hysteresis properties of titanomagnetites: grain-size and compositional dependence, Phys. Earth Planet. Inter. 13, 1977. [lo] J.-P. Valet and L. Meynadier, Geomagnetic field intensity and reversals during the past four million years, Nature 366, 91-95, 1993. [ll] L. Meynadier, J.-P. Valet, F. Bassinot and N.J. Shackleton, Assymetrical saw-tooth pattern of the geomagnetic field intensity from equatorial sediments in the Pacific and the Indian Oceans, Earth Planet. Sci. Lett. 126, 1099127, 1994. [12] J.-P. Valet, L. Meynadier, F. Bassinot and F. Gamier, Relative paleointensity across the last geomagnetic reversal from sediments of the Atlantic, Indian and Pacific Oceans, Geophys. Res. Lett. 21(6), 485-488, 1994. [13] W. Ruddiman, M. Samthein et al., Proc. ODP, Init. Rep. (Pt. A) 108, 1992. [14] J.-P. Pozzi and J. Zlomickl, Les variations d’aimantation des
[15]
[16]
[17]
[18]
[ 191
[20]
[21]
roches et le probleme des contraintes dans la croute terrestre, Phys. Earth Planet. Inter. 24, 169-177, 1981. N.J. Shackelton, S. Crowhurst, T. Hagelberg, N.G. Pisias and D.A. Schneider, A new late Neogene time scale: application to Leg 138 sites, Proc. ODP, Sci. Results 138, 1994. J.E.T. Channell, R. Freeman, F. Heller and W. Lowrie, Timing of diagenetic haematite growth in red pelagic limestones from Gubbio (Italy), Earth Planet. Sci. Len. 58, 189-201, 1982. K.A. Hoffman and S.B. Slade, Polarity transition records and the acquisition of remanence: a cautionary note, Geophys. Res. Lett. 13, 483-486, 1986 A.A.M. van Hoff and C.G. Langereis, The Upper Kaena sedimentary geomagnetic reversal record from southern Sicily, J. Geophys. Res. 98, 7939-7957, 1992. M. Hyodo, Possibility of reconstruction of the past geomagnetic field from homogeneous sediments, J. Geomagn. Geoelectr. 36, 45-62, 1984. T. Yamazaki, Thickness of the lock-in zone of the post-depositional remanent magnetization in deep-sea siliceous clay, Rock Magn. Paleogeophys. 11, 85-90, 1984. P.B. de Menocal, F. Ruddiman and D.V. Kent, Depth of post-depositional remanence acquisition in deep-sea sediments: a case study of the Brunhes-Matuyama reversal and oxygen isotopic stage 19.1, Earth Planet. Sci. Lett. 99, 1-13, 1990.
Earth and Planetary Science Letters 136 (1995) 541-550
Continuous record of geomagnetic field intensity between 4.7 and 2.7 Ma from downhole measurements J. Thibal a, J.-P. Pozzi a, V. Barth&s b, G. Dubuisson a a EC& Normale Sup&rieure. Luboraroire de Giologie, URA 1316 CNRS, 24 rue Lhomond, 75231 Paris Cedex 05, France h Centre Etudes Nu&aires, Laboratoire d’Electronique de Technologie et Instrumentation, 38041 Grenoble, France
Received 30 May 1995; revised 30 August 1995; accepted 1 September 1995
Abstract A continuous record of the geomagnetic field intensity from 4.7 to 2.7 Ma has been obtained, together with a precise magnetostratigraphy, from downhole magnetic measurements at Site 884 of ODP Leg 145 in the North Pacific. The record confirms the saw-tooth pattern of geomagnetic field intensity proposed by Valet et Meynadier [lo]. Reversals are characterized by a steep intensity decrease followed by a quick regeneration. Over each polarity interval, rapid variations are superimposed over a progressive decrease in the mean intensity. We find that the duration of each polarity interval is inversely proportional to the mean rate of decrease in the field intensity over this period, and thus this duration seems to be pre-determined.
1. Introduction Five years ago, LET1 (French Atomic Agency, CEA), CFP TOTAL and ENS-CNRS initiated a programme with the help of the Schlumberger Company to identify magnetostratigraphy in sediments by using in-situ downhole measurements. The feasability of in-situ magnetostratigraphy was first demonstrated within the weakly magnetized sediments of the Paris Basin [ll. We used a total magnetic field tool with a nuclear resonance magnetometer. To remove the effects of geomagnetic variations, a reference magnetometer was set up far from the disturbance of the rig (at sea, the nearest permanent landbased magnetic observatory is used). The induction measured within the borehole results from both induced and remanent magnetization of the rocks. A
susceptibility log obtained with a dipole-dipole susceptibility tool is then recorded (both tools are now combined). The processing of the combined induction and susceptibility data gives the polarity of the remanence. The induction B(z) generated in the borehole (as a function of depth) by magnetization J(z) is given by Bii( z) = Cij . Ji( z>, where C is a secondorder symmetrical tensor [Z-4] and J(z) is the total magnetization expressed by J(z) = Ji( z) + J,(z) (where Ji and J, are the induced and remanent magnetization). To a first order of approximation, the nuclear resonance magnetometer continuously measures the projection B’(z) of B(z) in the direction of the present-day geomagnetic field. Part B:(z) of B’(z), which is due to the remanent magnetization J,(z), reverses during reversals while Bi, which
0012-821X/9.5/$09.50 0 1995 Elsevier Science B.V. All rights reserved SSDI 0012-821X(95)00173-5
J. Thibal et al./Earth
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and Planetary Science Letters 136 (1995) 541-550
2. Downhole magnetostratigraphy
is due to the induced magnetization Ji( z> = x( z> . F, does not (x(z) is the susceptibility and F is the intensity of the present-day geomagnetic field). Bi( z) is calculated from the susceptibility log and is then subtracted from B’(z) to obtain B:(z) and the polarity of J,(z). J,(z), the natural remanent magnetization, can include viscous magnetization and remagnetizations. In these conditions and for a geological interval without large directional changes in J,(z), B:(z) is proportional to the intensity of J,(z). During ODP Leg 145, in the Aleutian Sea, magnetic logging was run at several sites. The scientific goal of the leg was a study of the paleoceanography and paleoclimatology of the North Pacific Ocean for the past 10 m.y. Fig. 1 shows the location of the sites drilled, and especially the deepest of the three sites (884) situated on the northern extremity of the Emperor Seamounts, on the Meiji Tongue [5]. In this paper we will show that the geomagnetic field paleointensity can be obtained from the magnetic logging. We present the geomagnetic field intensity variations together with the precise magnetostratigraphy obtained in the same hole (884E), situated at 51”27.034’N, 168”20.216’E. The samples recovered within a hole drilled a few metres away have been used to control the homogeneity of the magnetic mineralogy.
The magnetostratigraphic interpretation at hole 884E 161 is shown in Fig. 2. The susceptibility log (Fig. 2a) and the magnetometer record corrected for the pipe effect (Fig. 2b) are interpreted in terms of a polarity sequence (Fig. 2~). A comparison between this sequence and the geomagnetic polarity timescale from Cande and Kent [7] shows a perfect agreement down to 450 m below the seafloor (bsf), which corresponds to an age of more than 8 Ma. Consequently, the drilled formations down to 450 m bsf and the magnetic logs are dated. There is good agreement between the interpretation of logs for Hole 884E (Fig. 2c) and the polarities deduced from the core mesurements for hole 884B (Fig. 3).
3. Magnetization
and geomagnetic
paleointensity
Paleointensity studies rely on the assumption that the natural remanent magnetization (NRM) of sediments is mainly due to particle alignment under the effect of the torque exerted by the paleogeomagnetic field Fp during the sedimentation process. The intensity of the natural remanent magnetization J, can be expressed by: J, (A/m)
(
= A . C, . Fp (A/m)
(1)
1
.z,.:”8851886 North Pacific Ocean
I
I
1SO'E
**
;, A
0
160"
0 ,,.
.“i,_,
,
.:
I 150”
I
Fig. 1. Location map of the North Pacific Ocean showing Leg 145 drill sites.
1200
J. Thibal et al. / Earth and Planetary Science Letters 136 (I 995) 541-550
B’(z) NT) -0
Geomagnetic polarity time scale
Downhole polarity sequence
where A is a dimensionless coefficient which depends on the nature of the magnetic minerals, and on the size and shape of the grains of these minerals. C, is the concentration of ferromagnetic minerals. Provided A is constant over the considered logged section, one can normalize the concentration C, of
e 0
543
Polarity Inclination
1
-90”
0
900
sequence from cores
0
50
Downhole
polarity sequence 3
100
;,
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200
s 8 4.
150
250. 3
200
,I< .x_,.
.% 300 6
350
I’
I
a 8 a
250
.%
i 250
a & n
300
400
8
Fig. 2. Magnetostratigraphic results at Hole 884E. (a) Susceptibility log from dipole-dipole probe below the casing. (b) Borehole induction from the nuclear resonance magnetometer due to both remanent and induced magnetization of sediments (nT). (c) Interpretation as a downhole polarity sequence. Black/white = normal/reverse polarities. (d) Comparison with the geomagnetic polarity timescale of Cande and Kent [7].
350
400
Fig. 3. Comparison between the polarity sequence deduced from the inclination of the remanent magnetization measured on cores at Hole 884B and the downhole polarity sequence.
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J. Thibal et al. / Earth and Planetav
magnetic minerals. We limited our first study to the interval 4.7-2.7 m.y. (259-128 m bsf). After 2.7 m.y., the magnetic susceptibility displays large variations (Fig. 21, due to ash and dropstone layers becoming frequent in the sedimentation as a consequence of large climatic changes [5].
4. Rock magnetism experiments The following rock magnetism experiments on 70 samples demonstrate that the magnetic mineralogy is homogeneous over the selected section and that the concentration of magnetic minerals is well repre-
Science Letters 136 (1995) 541-550
sented by the variations in the low field susceptibility. 4.1. Magnetic minerals and grain size Cycles of hysteresis have been conducted on seventeen samples up to 6 T and at two temperatures (- 83°C and 20°C respectively), below and above Morin’s transition characteristics of haematite. The results are identical, demonstrating that haematite does not exist in detectable amounts and cannot be detected by Morin’s transition in this type of oceanic sediment. A typical result is shown in Fig, 4 (sample 164). Fig. 4a shows the - 0.1 to 1 T part of the cycle
Sample 164
Sample 225
PlOO
3
6 2
50
.;
para- 9.209 E-05
i
z
0
+ I.-0.2
0
0.2
0.4
0.6
o.*
1
Field (T)
z 2
50
f .z R ‘g 40
------------__--
6 Js- 50.17
E 0 ‘Z 30 6
Hc-
E g g 20 -1
1
cl
2 3 Field (T)
,,,,
A/m
11.4mT
0
5
6 Field (T)
““,.“‘,““,
2 2:
JrsHrc= 8.30 31.7 A/m mT
: Hrc= 46.4
mT
0 -2 -4 -0.5
+ ..,.l~,~~l~.~.l.,..
.‘I.
0
0.5
1
Field (T)
1.5
2
-0.5
0
0.5
1
1.5
2
Field (T)
Fig. 4. Two typical examples of high field measurements. Sample 164 is typical of the 131 m of the section, except for the interval 180.5-196 m bsf. Sample 225 is typical of interval 180.5-196 m bsf, showing a large proportion of a haematite-like mineral, and is not considered further.
J. Thibal e? al. / Earth and Planetary Science Letters 136 (1995) 541-550
0.5
Hole
.‘.’ ,
0.61,
L I SD 1
I
0.4 1
I
“.
8848
I”’
““I’
‘2 : 7
PSD !
1
H,/H,.
This diagram to obtain an estimate The results are given grain size is quite single-domain (PSD) 4.2. Concentration
L
ol ,, 0
!,,I, 1
~,,!?,I
//,1
I,,,,’
2
3
4
5
6
HrclHc
Fig. 5. Grain-size diagram deduced from hysteresis properties. J,, /J, vs. H,, /H, [9]. The fourteen samples taken over the considered section are confined in the pseudo-single-domain (PSD) range.
which has been conducted up to 6 T. Fig. 4b illustrates the ferromagnetic part of the cycle when the paramagnetic contribution is subtracted. The saturation magnetization J, is obtained for 1.8 T and the coercive force H, is found to be 11.4 mT. Fig. 4c gives the saturation isothermal remanence magnetization (IRM) 1, and the coercive force of the remanence H,. The saturation of IRM is reached at 0.3 T. For all specimens (fourteen) represented by this typical curve, H, ranges from 10.9 to 26.2 mT with a mean value of 14.6 mT and H, ranges between 31.7 and 80.2 mT with a mean value of 43.1 mT. These measurements show that fine-grained magnetite is bound to be the main magnetic mineral. Nevertheless, in the defined depth interval 196180.5 m bsf, the susceptibility log (Fig. 2a) displays very low values, in most cases lower than the minimum value for the probe pre-set for the particular logging operation in question. Fig. 4 shows the high field measurements of sample 225, which is representative of the specimens sampled in this zone. Fig. 4d-f refers to the same measurements as Fig. 4a-c. The results on samples from this interval show the occurrence of a high-coercivity mineral, probably haematite. The fact that no difference has been detected in the hysteresis parameters obtained at the two temperatures above and below Morin’s transition may be due to a content of very fine grained haematite [8]. The grain size of the magnetic mineral is determined from the classical diagram for J,,/J, vs.
545
is well adapted in this case [9] of the grain-size homogeneity. in Fig. 5, and indicate that the homogeneous in the pseudorange.
Cr and susceptibility
Anhysteretic remanent magnetization (ARM) is currently used in paleointensity studies to normalize J, to the variation in the ferromagnetic mineral concentration Cr. We first compare the alternating field demagnetization curves for NRM and ARM. Typical results are given in Fig. 6 and show that the two demagnetization curves are similar. Then we compare the ARM and low field susceptibility x. The crossplot of the two parameters (Fig. 7) indicates that the relationship between susceptibility and ARM is close to proportionality over the section. Therefore, we can use the downhole susceptibility x to normalize the variations in Cr. The above considerations lead us to propose that the paleointensity variations in the geomagnetic field can be obtained from downhole magnetic measurements: from the total field log B’(z) we subtract the induced part B:(z), and so we obtain a signal for B$ z) proportional to J,(z). The effect of the variation in the ferromagnetic mineral concentration C, is normalized to a constant concentration using B[( z).
301,"
"
',"I/,88
”
,,,T
Sample
174
i
300
r] 250
j
gc 1DL 5
L :
o
0
100
SNNRM L
z Q
50 --ARM
.,,','I 10
""'l.,,,i..,.!,,,,,o 20
30
40
50
60-
Field (mT) Fig. 6. Typical plot showing alternating field demagnetization of both natural remanent magnetization (NRM) and anhysteretic remanent magnetization (ARM; 100 mT, 0.1 mT). The similarity of the two curves shows that the same magnetic grains are involved in the two processes.
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and Planetary Science Letters 136 (1995) 541-550
Thus, the variations in the geomagnetic field paleointensity can be expressed as [B’(z) - BI(z)]/B:(z), which can be obtained from the magnetic logs.
5. Paleointensity record
03 0
100
200
300
400
500
600
700
600
Susceptibility (~10~ SI) Fig. 7. ARM after a 20 mT alternating field demagnetization vs. low field susceptibility x for single samples. The linear relationship between the two parameters indicates that the low field susceptibility is a suitable normalizing parameter for the NRM. It also confirms a uniform magnetic granulomety over the considered section.
2.7
3.3
3.5
The resulting relative paleointensity record is given in Fig. 8, with positive values for normal polarity (black) and negative values for reverse polarity (white). The sampling depth interval of the magnetic tool is 0.15 m, which corresponds to a maximum of 3.5 k.y. and a minimum of 1.5 k.y. respectively for the lowest and the highest values of the accumulation rate deduced from the magnetostratigraphy over the selected section (43 m 1 m.y.-’ and 96 m . m.y. - ‘1. The NRM/ARM ratios measured on specimens after alternating field demagnetization up to 20 mT are shown in the same figure for comparison. There is a good agreement
3.7
3.9
4.1
4.3
4.5
4.7
Age (MY) Fig. 8. Relative variation of the geomagnetic field intensity over the 4.7-2.7 Ma interval obtained from downhole magnetic records at ODP Site 884. The graph is shown in a polarized form (positive and negative values for normal (black) and reverse (white) polarity periods). 0 = NRM/ARM ratios mesured on specimens after AF demagnetization up to 20 mT.
J. Thibal et al./Earth
and Planetary Science Letters 136 (1995) 541-550
4.7-2.7 Ma, confirm the saw-tooth pattern presented by Valet and Meynadier [lo] and extend the record of continuous paleointensity 0.7 m.y. beyond their record. In addition, Valet and Meynadier’s record is representative of field behaviour at low, equatorial latitudes, whereas our result is the first obtained at high latitude. Note that from Valet and Meynadier’s data, the mean virtual axial dipole moment is 4.1 f 2.0 X lo** A m* for normal periods and 3.3 f 1.5 X lo’* A m* for reverse periods. The mean normal and reverse paleointensities calculated for 4.7-2.7 Ma at Site 884 are equal to within a few percent. In both our and Valet and Meynadier’s study the same paleointensity method is applied (although on cores by Valet and Meynadier and on the surrounding sediments within the borehole in our case). Cores can be remagnetized in the core barrel assembly [ 131. A further remagnetization is created during storage of the core while minerals reach their new physical and chemical equilibrium. In addition, coring certainly remagnetizes the wall of the borehole by the effect of the field produced by the drilling and coring device, probably enhanced by stress relaxation and piezomagnetic effects [14]. However, the remagnetized volume of the surrounding rock is very small compared to the tool investigation volume, and the remagnetized volume is, therefore, relatively much larger in the core. Yet, the comparison between the two records shows that re-
between laboratory and downhole measurements. This comparison validates our method based on rapid data acquisition. It has the advantage of not being dependent on core recovery and does not require demagnetization. The paleointensity curve is interrupted over the two ash-dominated intervals from 4.12 to 4.10 m.y. and from 3.50 to 3.49 m.y. Between 3.75 and 3.56 m.y. (respectively 196 and 180.5 m bsf) a highcoercivity haematite-like mineral is dominant, whereas elsewhere magnetite is dominant. We therefore suspect that the coefficient A (Eq. 1) changes. We have no explanation for this localized change in the proportion of magnetic minerals, except that it corresponds to a sharp decrease in clay content [5] which could have enhanced the permeability and favoured fluid circulation. Several characteristic features of the geomagnetic field can be shown on the curve. During the interval of constant polarity, the behaviour of the field intensity is characterized by large variations superimposed on a steady decrease. Each reversal seems to be the result of the coupling of the rapid variations with the progressive decrease in the mean intensity and is followed by a rapid regeneration of the field strength. These features confer on the curve an asymmetrical saw-tooth shape. The only available previous study [lo-121 has shown a synthetic record obtained from several equatorial Pacific and Indian Ocean cores. Our results, which span the period
2.7
2.8
2.9
3
3.1
3.2
54-l
3.3
3.4
3.5
3.6
3.7
3.8
3.9
4
Age(My)
Fig. 9. Comparison between Valet and Meynadier’s relative paleointensity curve established from core measurements from the Indian and Pacific Oceans at low latitude and our curve (heavy line) deduced from downhole magnetic records at ODP Site 884 at high latitude for the interval 4-2.7 Ma. Curves are plotted with a common age scale and a positive constant is added to our paleointensity.
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J. Thibal et al./ Earth and Planetary Science Letters 136 (1995) 541-550
magnetization probably does not affect their quality because the cores have been demagnetized in the laboratory. Fig. 9 shows the detailed comparison of the two records along the common period 4.7-2.7 Ma. To make the comparison with Valet and Meynadier’s results easier, we have tuned our record to the same age scale [15] and a positive constant has been added to the relative paleointensity. The similarity in the records is excellent (even in detail), and very satisfying correlations can be established. Note, however, that slight discrepancies in age within constant polarity periods cannot be avoided. These result from specific time distortion of each record due to rapid variations in accumulation rate that cannot be uncovered by magnetostratigraphy. In addition, let us also consider the effect of a lock-in depth of natural remanence in the paleofield-remanence relationship. Such an aquisition depth of post-depositional remanent magnetization constitutes a zone that can be seen as the result of a mixing of normal and reverse grains just below a reversal [ 16-181, but also as a mixing of grains magnetized to varying intensity throughout the polarity periods between reversals. This it affects the signature of intensity variations during periods of contant polarity. This lock-in zone, which is assumed to have a constant thickness, represents a time dura-
2.7
2.9
3.1
3.3
3.5
tion which depends on the accumulation rate. The zone acts on the sedimentary record of paleointensity as a high-frequency filter whose cut-off frequency is proportional to the accumulation rate. The effect of this zone must not be underestimated. For the lowest accumulation rate at Site 884, a lock-in zone of 0.15-0.30 m in thickness as currently estimated in deep-sea sediments [19-211 has a duration of 7 k.y. A detailed study of the effect of the lock-in zone, which would require comparison with other paleointensity records in holes with different sedimentation rates, is unfortunately beyond the scope of this paper. We next study polarity periods excluding transitions themselves, and analyse the decrease in mean field intensity during each period, without attempting to describe in detail the rapid and large variations superimposed over this decrease. Fig. 10 shows the same data as Fig. 8 in an unpolarized form (i.e., not taking polarity sign into account). Valet and Meynadier [lo] suggested that the polarity chron length is related to the regeneration of the field intensity following a transition. In other words, polarity length is proportional to the field intensity jump across the transition. However, some bias can arise from the difference between the mean dipole field intensity of the normal and reverse polarities.
3.7
3.9
4.1
4.3
4.5
4.r
Age (MY) Fig. 10. The relative variations form. A linear fit is calculated
of the geomagnetic field intensity over the 4.7-2.7 Ma interval at ODP Site 884 are plotted in an unpolarized within each polarity Period between 4.7 and 2.7 Ma after removing transitions.
J. Thibal et al. / Earth and Planetary Science Let&n
.L
6
80
@ f; cl
E
0
L
I
I
0.1
I
0.2
1
1
0.3
0.4
0.5
Duration of polarity intervals P (MY) Fig. 11. Inverse linear relationship between the rate of decrease of the mean relative paleointensity over a polarity period and the duration of the corresponding period. R is expressed in percent per 0.1 m.y. of the average intensity calculated over the 4.7-2.7 Ma interval. The error bars on the duration of polarity intervals, ranging between 7 and 13%, are deduced from errors in the timescale assumed by Cande and Kent from the duration of chrons [7]. The R error bars are proportional to the residuals of the linear approximation.
Our record leads us to propose another relationship. We have calculated the mean rate R of decrease in the intensity over each polarity interval, using a linear curve fitting (the only one justified by the data), after removing the transitional data within a 40 k.y. window centred qn the reversal (Fig. 10). It seems that the length of a polarity period P is related to the linear rate of decrease R in the field intensity. The relationship between P and R is shown in Fig. 11, which shows that P is inversely proportional to R.
6. Conclusion This case study confirms the possibility of obtaining a precise magnetostratigraphy from continuous downhole magnetic measurements [l]. Our main new result is that a continuous paleointensity record of the Earth’s magnetic field can be obtained from a single logging operation in a deep-sea sedimentary sequence. This was successfully done at ODP Site 884 in the North Pacific. This new record of paleointensity, obtained for a high latitude, is in excellent
I36 {I 995) 541-550
549
agreement with the record of Meynadier et al. [l 11 deduced from core mesurements at low latitudes. Each reversal seems to be the result of the coupling of the rapid variations with the progressive decrease of the mean intensity. Reversals are characterized by an intensity decrease followed by a rapid regeneration of the field strength. The saw-tooth pattern is confirmed and extended by 0.7 m.y. During periods of constant polarity, the paleointensity shows rapid variations superimposed on a slow decrease in the mean field. We also propose a new law for the field intensity behaviour between reversals: It seems that the length of a polarity interval is inversely proportional to the mean rate of decrease in the field intensity during this interval. This leads to the surprising result that the duration of a given polarity interval could be immediately determined after the previous reversal.
Acknowledgements We thank X. Le Pichon and J.-P. Valet for helpful comments and J. Pocachard and A. Etchecopar for their contribution to this programme. The high field measurements were conducted with the help of N. Bontemps and P. Monod in the Physics Department of the ENS using a Squid magnetometer. The data used for this study are available on request from the BRG of the LDEO. This work was partly supported by Schlumberger-EPS and by CNRS-INSU (Programme Giosciences Marines). [ PT]
References 111J.-P. Pozzi, V. Barthes, J. Thibal, J. Pocachard,
M. Lim, T. Thomas and G. Pages, Downhole magnetostratigraphy in sediments: comparison with the paleomagnetism of a core, J. Geophys. Res. 98, 7939-7957, 1993. 121J.-P. Pozzi, J.-P. Martin, J. Pocachard, H. Feinberg and A. Galdeano, In-situ magnetostratigraphy: interpretation of magnetic logging in sediments, EarthPlanet. Sci. Len. 88, 357373, 1989. in [31 Y. Gallet and V. Courtillot, Modeling magnetosuatigraphy a borehole, Geophysics 54, 973-983, 1989. [41 Y. Hamano and H. Kinoshita, Magnetization of the oceanic crust inferred from magnetic logging in Hole 395A, Proc. ODP, Sci. Results 106-109, 1990. 151 D.K. Rea, LA. Basov et al., Proc. ODP, Init. Rep. 145, 1992.
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[6] G. Dubuisson, J. Thibal, V. Barthes, J. Pocachard and J.-P. Pozzi, Downhole magnetic logging in sediments during Leg 145: usefulness and magnetostratigraphic interpretation of the logs at site 884, Proc. ODP, Sci. Results 145, in press. [7] S.C. Cande and D.V. Kent, A new geomagnetic polarity time scale for the late Cretaceous and Cenozoic, J. Geophys. Res. 97, 13917-13951, 1992. [8] F.D. Stacey and SK. Banerjee, The Physical Principles of Rock Magnetism 5, Elsevier, New York, 1974. [9] R. Day, M. Fuller and V.A. Schmidt, Hysteresis properties of titanomagnetites: grain-size and compositional dependence, Phys. Earth Planet. Inter. 13, 1977. [lo] J.-P. Valet and L. Meynadier, Geomagnetic field intensity and reversals during the past four million years, Nature 366, 91-95, 1993. [ll] L. Meynadier, J.-P. Valet, F. Bassinot and N.J. Shackleton, Assymetrical saw-tooth pattern of the geomagnetic field intensity from equatorial sediments in the Pacific and the Indian Oceans, Earth Planet. Sci. Lett. 126, 1099127, 1994. [12] J.-P. Valet, L. Meynadier, F. Bassinot and F. Gamier, Relative paleointensity across the last geomagnetic reversal from sediments of the Atlantic, Indian and Pacific Oceans, Geophys. Res. Lett. 21(6), 485-488, 1994. [13] W. Ruddiman, M. Samthein et al., Proc. ODP, Init. Rep. (Pt. A) 108, 1992. [14] J.-P. Pozzi and J. Zlomickl, Les variations d’aimantation des
[15]
[16]
[17]
[18]
[ 191
[20]
[21]
roches et le probleme des contraintes dans la croute terrestre, Phys. Earth Planet. Inter. 24, 169-177, 1981. N.J. Shackelton, S. Crowhurst, T. Hagelberg, N.G. Pisias and D.A. Schneider, A new late Neogene time scale: application to Leg 138 sites, Proc. ODP, Sci. Results 138, 1994. J.E.T. Channell, R. Freeman, F. Heller and W. Lowrie, Timing of diagenetic haematite growth in red pelagic limestones from Gubbio (Italy), Earth Planet. Sci. Len. 58, 189-201, 1982. K.A. Hoffman and S.B. Slade, Polarity transition records and the acquisition of remanence: a cautionary note, Geophys. Res. Lett. 13, 483-486, 1986 A.A.M. van Hoff and C.G. Langereis, The Upper Kaena sedimentary geomagnetic reversal record from southern Sicily, J. Geophys. Res. 98, 7939-7957, 1992. M. Hyodo, Possibility of reconstruction of the past geomagnetic field from homogeneous sediments, J. Geomagn. Geoelectr. 36, 45-62, 1984. T. Yamazaki, Thickness of the lock-in zone of the post-depositional remanent magnetization in deep-sea siliceous clay, Rock Magn. Paleogeophys. 11, 85-90, 1984. P.B. de Menocal, F. Ruddiman and D.V. Kent, Depth of post-depositional remanence acquisition in deep-sea sediments: a case study of the Brunhes-Matuyama reversal and oxygen isotopic stage 19.1, Earth Planet. Sci. Lett. 99, 1-13, 1990.