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Magnetostratigraphy from downhole measurements in ODP holes
Physics of the Earth and Planetary Interiors 156 (2006) 261–273
Magnetostratigraphy from downhole measurements in ODP holes Trevor Williams ∗ Borehole Research Group, Lamont-Doherty Earth Observatory, 61 Route 9W, Palisades, NY 10964, United States Received 24 May 2004; received in revised form 14 July 2005; accepted 22 September 2005
Abstract The “Geological High-resolution Magnetic Tool” (GHMT) was used on 17 ODP Legs to obtain downhole logs of magnetic field strength and magnetic susceptibility. The remanent magnetization of the sediment around the borehole is derived from these measurements, and used to produce magnetostratigraphies. This review of the ODP GHMT data indicates that reliable magnetostratigraphies can usually be obtained from reasonably well-magnetized (above approximately 10−3 A/m) fine-grained sediments. Additionally, the GHMT data indicates that a drill-string overprint, common in data collected on cores, does not significantly affect the sediments surrounding the borehole. This allows us to use the remanent and induced magnetization data to obtain estimates of relative paleointensity, and when averaged over time in homogenous sediment, the paleomagnetic recording efficiency of that sediment. © 2006 Elsevier B.V. All rights reserved. Keywords: Magnetostratigraphy; ODP; Downhole measurements; Paleomagnetism; Paleointensity
1. Introduction The “Geological High-resolution Magnetic Tool” (GHMT) was developed as a cooperative venture between Commissariat a l’Energie Atomique–Laboratoire d’Electronique, de Technologie et d’Instrumentation (CEA–LETI), the Ecole Normale Superieur, and the Total Oil Company in France, and has been operated in the Ocean Drilling Program (ODP) by Schlumberger starting in 1990 (ODP Leg 134). The motivation for developing the tool was to provide magnetostratigraphic dating of formations that had only limited biostratigraphic dating from either spot coring or cuttings. Magnetostratigraphic dating requires good time-sampling
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of the stratigraphic section, and the GHMT provides this by coverage of the whole open hole interval. In conventional paleomagnetic measurements on cores, Uchannels, and samples, several checks on the quality of the data are available, such as sequential demagnetization and visual inspection of the sediment for deformation or inhomogeneities. Additionally, paleomagnetic measurements are made in low-field environments, so there is no induced component to the sample’s magnetization. In measurements made downhole, demagnetization and visual checks are not possible, and the sediment has induced and viscous magnetizations in addition to the remanent magnetization which carries the magnetostratigraphy. Despite the differences, these concerns can be assessed: questions about the GHMT data and its interpretation are addressed in this review. Compilation of the ODP GHMT data set and comparison to paleomagnetism measured on cores recovered from the
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same holes indicates that valid magnetostratigraphies are obtained by the GHMT in reasonably well-magnetized sediment sections, for example down to approximately 10−3 A/m. Given a strong magnetization, the GHMT can provide information about relative paleointensity and the paleomagnetic recording efficiency of different sediment types. The GHMT measures the magnetic susceptibility in the formation surrounding the borehole using induction, and measures the total magnetic field in the borehole using a proton-precession magnetometer. In order to access the magnetostratigraphy, the remanent field anomaly must be isolated from the total field measured in the borehole. The measured field is dominated by the Earth’s dipole field, but is also contributed to by local fields due to the drill pipe, basement rocks, inhomogeneities in the formation, and the induced magnetization of the sediment. All of these must be subtracted to leave the remanent field anomaly: the quality of the magnetostratigraphy depends on how large the remanent anomaly is compared to the errors and noise in the estimations of the other fields. The theory describing the magnetic field in a borehole caused by the magnetization of the rocks surrounding the borehole was described by Parker and Daniell (1979) to enable interpretation of future downhole magnetometer observations in terms of the heterogeneity of magnetization in basaltic ocean crust. Pozzi et al. (1988) and Gallet and Courtillot (1989) developed the theory for horizontal and dipping magnetized layers with the aim of deriving the magnetostratigraphy of sedimentary successions. Initial field tests of the GHMT are described for the Permian Lod`eve Basin in the south of France (Pozzi et al., 1988), and for the Paris Basin (Bouisset and Augustin, 1993; Pozzi et al., 1993). GHMT results for ODP holes are presented for Leg 134 (Roperch et al., 1994), Leg 145 (Dubuisson et al., 1995), Leg 162 (Higgins et al., 1999), Leg 165 (Louvel and Galbrun, 2000), and Leg 178 (Williams et al., 2002). The GHMT was also run on ODP Legs 155, 160, 171B, 175, 184, 188, and 189, as summarized in Table 1. GHMT magnetic susceptibility records were obtained on ODP Legs 154, 167, 177, 181, and 182. Besides the GHMT proton-precession magnetometer, a variety of downhole fluxgate magnetometers have been used in ODP and DSDP boreholes for tool orientation or investigation of the magnetization of igneous ocean crust. Schlumberger’s general purpose inclinometry tool (GPIT) is a three-axis fluxgate tool that has been used on most ODP legs since Leg 126 mainly for orientation of downhole resistivity images, but also occasionally for paleomagnetism, for example in oceanic
basement in Hole 504B (Kinoshita et al., 1989) and in Hole 801C (Ito et al., 1995). A Russian three-axis fluxgate magnetometer was first used on Legs 68 and 69, and successfully detected reversals between basalt flows in Hole 395A on DSDP Leg 78B (Ponomarev and Nechoroshkov, 1983, 1984). A Japanese three-axis fluxgate magnetometer was also run in Hole 395A on Leg 109 (Hamano and Kinoshita, 1990). The Japanese magnetometer was also run on Leg 143 in both the sediments of Hole 869 (Nogi et al., 1995a) and the basalts of Holes 865 and 866 (Nogi et al., 1995b). A gyroscopeoriented three-axis fluxgate magnetometer, developed in Hannover, Germany, was run in Hole 418A, Leg 102, yielding magnetic characteristics of basalts and paleomagnetic pole positions (Bosum and Scott, 1988). Pariso et al. (1991), compared results from two downhole fluxgate magnetometers, one from the University of Washington and one from the US Geological Survey, in the Gabbros of Hole 735B. More recently, the Gottingen Borehole Magnetometer was run in alternating basalts and volcaniclastic sediments in Hole 1203A of Leg 197 (Tarduno et al., 2002). It comprises a three-axis fluxgate magnetometer together with an optical angular rate sensor, which, in theory, is accurate enough to determine the orientation of the tool independently from the magnetic reference frame, enabling the horizontal direction of the field to be determined. The advantage of the proton precession magnetometer over the fluxgate tools is its accuracy (0.1 nT), relative insensitivity to temperature variations, and its lack of measurement drift. The three-axis fluxgate magnetometers provide the vertical and horizontal components of the field, but the absence of directional information in the total field measurement of the GHMT is not generally a barrier to polarity determination, as discussed below. The GHMT magnetic susceptibility logs provide a valuable lithological indicator sensitive to paleoceanographic and paleoclimatic change, as well as a good parameter for correlation between cores and downhole logs. This paper is concerned solely with the paleomagnetic applications of the GHMT, and the paleoclimatic applications are not discussed further here. 2. Derivation of polarity of remanent magnetization from the total magnetic field measured in a borehole 2.1. Remanent magnetization The remanent magnetization (Jr ) of sediments is typically acquired during deposition by preferential alignment of the magnetic grains along the magnetic field
Table 1 Magnetic field and magnetic susceptibility logs were obtained by the GHMT in the holes listed in this table Leg
Hole
Location
Depth range (m)
Age
Induced anomaly range (nT)
1990 1992 1992
134 145 145
831B 883F 884E
Bougainville Guyot Detroit Seamount Detroit Seamount
100–700 220–720 90–680
Oligocene–Pleistocene Oligocene–l. Pliocene l. Miocene–l. Pliocene
30 70
1994 1994 1995 1995 1995
155 155 160 160 162
931B 933A 966F 967E 986C
Amazon Fan Amazon Fan Eratosthenes Seamount Eratosthenes Seamount Svalbard Margin
90–240 90–210 100–340 100–580 100–360
l. Pleistocene l. Pleistocene Eocene–l. Pliocene l. Pliocene Pleistocene
1995
162
987E
Greenland Margin
100–300
1996 1996 1997 1997 1997 1997 1997 1997 1998
165 165 171B 171B 171B 175 175 175 178
998B 1001A 1050C 1051A 1052E 1081A 1082A 1084A 1095B
Cayman Rise Nicaraguan Rise Blake Nose Blake Nose Blake Nose Walvis Ridge Walvis Basin Cape Basin Antarctic Peninsula
1998
178
1096C
1998 1999 1999 1999 2000 2000 2000 2000
178 184 184 184 188 189 189 189
1103A 1144A 1146A 1148A 1166A 1168A 1170D 1172D
Remanent anomaly range (nT)
Invalid data spikes
Reference
30 100
None None Present
Roperch et al. (1994) Dubuisson et al. (1995) Dubuisson et al. (1995), this paper
25 20 3 3 80
20 20 2 8 100
None Present None Present Present
l. Pliocene–Pleistocene
100
150
Present
210–430 200–400 125–585 140–590 240–680 100–400 90–550 80–590 110–550
Oligocene–e. Miocene Cretaceous–Paleocene Cretaceous–Eocene Cretaceous–Eocene Cretaceous–Paleocene l. Miocene–l. Pliocene l. Pliocene–e. Pleistocene l. Pliocene–l. Pleistocene l. Miocene–e. Pliocene
5 6 12 10 10 8 8 8 100
8 10 4 3 3 8 12 25 200
None None Present Present Present Present Present Present Present
Antarctic Peninsula
360–500
e. Pliocene
100
200
None
Antarctic Peninsula N Margin S China Sea N Margin S China Sea N Margin S China Sea Prydz Bay West Tasmania Slope South Tasman Rise East Tasman Plateau
100–240 120–440 250–585 400–705 40–380 140–720 550–780 210–760
l. Pliocene–l. Pleistocene l. Pleistocene m. Miocene–e. Pliocene Oligocene–e. Miocene Eocene-Pliocene e. Oligocene–l. Miocene Eocene Cretaceous–l. Miocene
350 6 6 5 300 35 35 50
100 6 6 2 50 10 10 10
Abundant Abundant Abundant Abundant Present Present None Present
Higgins et al. (1999), Channell et al. (1999) Higgins et al. (1999), Channell et al. (1999) Louvel and Galbrun (2000) Louvel and Galbrun (2000)
This paper This paper, Williams et al. (2002) This paper, Williams et al. (2002)
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Year
The amplitude ranges of the remanent and induced anomalies are approximate values for decameter scale variation. The number of invalid data spikes are classified here as none (if the spikes are absent or negligible), present (if spikes are present in all or part of the log and can be removed to leave good data), or abundant (if the spikes dominate the log and prevent analysis). GHMT-based magnetostratigraphies can generally be produced from holes where the remanent anomaly is equal to or higher than the induced anomaly. The GHMT data is available from http://iodp.ldeo.columbia.edu/DATA/index.html. 263
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direction. The remanent magnetization contributes to the field anomaly in the borehole (Bfr ), according to the following equations (Pozzi et al., 1988) (Fig. 1A): Bfrx =
µ 0
2
· Jrx
(1)
Bfrz = −µ0 · Jrz
(2)
where Bfrx and Bfrz are the components of Bfr in the x (∼north) and z (down) directions. The Earth’s main field, Br , is much larger than the field anomaly due to the remanent magnetization (in the example in Fig. 1, 44,000 nT compared to 114 nT), so the total field, B0 , and the Earth’s field, Br , are sub-parallel (the angle between the fields is 0.104◦ , Fig. 1B). Therefore, to a very close approximation, it is the projection of the remanent anomaly Bfrp along B0 that contributes to the measured B0 . This is the basis for determining polarity using the intensity of the total field rather than its directional change: whereas an ±80 nT change in total field is easy to measure, the directional change in the field is very difficult to measure, especially from a tool that is free to rotate and is subject to impacts in the borehole. Bfrp is calculated by (Pozzi et al., 1988): Bfrp =
µ0 · J r 2
· (1 − 3 sin2 I)
(3)
Fig. 2. Dependence on the field inclination of the remanent anomaly, Bfrp , in the borehole due to a 0.1 A/m remanent magnetization. Note that at an inclination of ±35◦ , no remanent anomaly is observed, because Bfr is perpendicular to the earth’s field (see Fig. 1B).
where I is the inclination of the remanent magnetization. Note that at an inclination of 35◦ , the field due to the remanent magnetization Bfrp is perpendicular to the total field, and the anomaly cannot be observed with total field measurement (Fig. 2). 2.2. Induced magnetization The sediment also carries an induced magnetization (Ji ) along the direction of the present-day field. Ji can be calculated from the magnetic induction H and the magnetic susceptibility κ (measured by the GHMT) (Eqs. (4)
Fig. 1. An example showing the remanent and induced magnetizations (Jr and Ji ) in the formation and the remanent and induced field anomalies they cause in the borehole (Bfr and Bfi ). Given a sediment magnetization, Jr , of 0.1 A/m, the observed anomaly in the borehole, Bfr , is 80 nT (see text). The example is from the southern hemisphere, at the latitude of Hole 1096B, where the field inclination is −60.6◦ . Note that the fields in the borehole do not have the same direction as the sediment magnetizations that produce them.
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and (5)). The field anomaly in the borehole Bfip caused by the induced magnetization Ji can then be calculated by using the equivalent of Eq. (3). B0 Ji = ·κ (4) µ0 B 0 = µ0 · H
(5)
where: B is the magnetic induction, unit T; H is the magnetic field, unit A/m; J is the magnetization, volume normalized, unit A/m; κ is the magnetic susceptibility, volume normalized, unitless (×10−6 SI units); µ0 is the permeability of nonmagnetic material = 4π × 10−7 (in SI units). 2.3. Deriving the GHMT polarity from the measured total field The measured total field is the vector sum of a number of different contributing fields, and the remanent anomaly is found by progressively subtracting the other individual fields from the total (see Fig. 3). B0 (z, t) = Br (z) + Ba (z) + Bfr (z) + Bfi (z) + Bt (z, t) (6) All the components of B0 are a function of depth, z; the time-dependant magnetic field, Bt , generated in the ionosphere, also contributes. Br , the field generated in the Earth’s core, forms the largest component of B0 by about two orders of magnitude. It is estimated from the International Geomagnetic Reference Field (IGRF). Ba is the regional magnetic field: sea-floor and other crustal anomalies. The field caused by the metal (magnetic) drill pipe (Bpipe ) is also included within Ba ; it decays with distance (r) away from the pipe, Bpipe = a/r3 . The pipe effect is determined by iterative forward modeling until a satisfactory fit to the data is obtained. Bt is the time dependant field, originating in the magnetosphere and ionosphere, which varies on timescales of minutes to days. Two passes of the GHMT are usually run so that the repeatability can be checked; for the most part the differences are slight and do not significantly affect the data. However, negative ‘spikes’ to low field values occur in many of the GHMT logs, and these spikes do not repeat from pass to pass: they originate in the tool itself and are not a valid recording of the field. They are accompanied by a drop in the tool voltage, and the invalid magnetic field data may be removed from the log on the basis of the voltage (Williams et al., 2002). Bfip , the induced field anomaly, is calculated using Eqs. (3)–(5), with values for B0 and I taken from the IGRF values for the site.
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Even after Br , Ba , and Bfip have been subtracted from the measured total field B0 to leave Bfrp , the baseline value of the remanent field anomaly, Bfrp can be offset from zero. This offset can be caused by diurnal variation, local discrepancies from the IGRF, proximity to basement crustal anomalies, etc. However, we know a priori that over enough normal and reverse polarity intervals Bfrp will average out to near zero. Therefore, we can correct for this baseline offset by subtracting a constant value or a linear fit, resulting in a remanent anomaly Bfrp that is positive in normal polarity sediment and negative in reversed polarity sediment. Thus, the first method for determining polarity is the sign of the remanent anomaly, which is often sufficiently good to provide a basis for magnetostratigraphy. A second method, correlation analysis, overcomes the problem of a variable baseline offset (Pozzi et al., 1993; Vibert-Charbonnel, 1996; Louvel and Galbrun, 2000). The principle employed by the correlation analysis method is illustrated in Fig. 4. Both Jr and Ji are dependant on the concentration of ferrimagnetic minerals in the sediment, and this concentration varies. Jr and Ji correlate in normal polarity intervals, and the linear regression line has a positive gradient, whereas in reversed polarity intervals Jr is inversely correlated with Ji and the linear regression line has a negative gradient. The same applies to Bfrp and Bfip , on which the correlation analyses are actually performed. The linear regressions are applied to successive depth intervals (windows) of various thickness (1.5, 2.5, 4.5, 8, and 13 m windows are presented in Fig. 3). Linear regressions with a correlation coefficient below 0.5 are not plotted. Prior to the correlation analyses, Bfrp and Bfip are smoothed with an 11 sample (1.5 m) Hanning filter, so that the Bfrp and Bfip logs have comparable vertical resolution. Results of the two methods are plotted side-by-side in Fig. 3: the sign of the remanent anomaly gives a more continuous polarity record, while correlation analysis gives a better reading close to the drill pipe. 2.4. Constraints on GHMT polarity determination In order to evaluate the qualitative confidence that should be applied to GHMT magnetostratigraphies, the assumptions involved in the analysis are discussed below, followed by presentation of examples of demonstrably successful GHMT magnetostratigraphies in ODP. The majority of the issues below affect only the amplitude of the remanent anomaly and not its polarity, therefore leaving the GHMT magnetostratigraphy minimally affected.
266 T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273 Fig. 3. An example from Hole 1096B showing the derivation of the remanent anomaly from the measured magnetic susceptibility and total field logs. The original measured logs are susceptibility and the total field. The induced anomaly is calculated from the susceptibility, and then the remanent anomaly is calculated by subtraction of the induced anomaly, the pipe effect, and a constant value from the total field. Polarity can be determined from (a) the sign of the remanent anomaly and from (b) the correlation analysis method (positive gradients in black, negative gradients in green). In each case a polarity column is given, with black indicating normal polarity, green reversed polarity, and white undetermined polarity. The depth window employed in the correlation analyses are (from left to right) 1.5, 2.5, 4.5, 8, and 13 m. Paleomagnetic data from Hole 1096B after demagnetization in peak fields of 20 mT and the GPTS of Cande and Kent (1995) are given for comparison.
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Fig. 4. Remanent (Bfrp ) vs. induced (Bfip ) anomalies for the interval from 400 to 420 mbsf in Hole 1096C, covering the Gilbert–Gauss polarity reversal, illustrating the principles of the correlation analysis method. Correlation between Bfrp and Bfip indicates normal polarity, and anti-correlation indicates reversed polarity.
2.4.1. Inclination of the remanent magnetization Both the amplitude and polarity of the remanent anomaly depend on the inclination of the remanent magnetization (Fig. 2), so the inclination must be known or estimated. In the simplest case of geologically young horizontally bedded sediments, the remanent magnetization will be either parallel or anti-parallel to the Earth’s axial dipole field. Most of the ODP boreholes logged with the GHMT approximate to this case. Paleomagnetic secular variation will generally average out to the axial dipole field over the volume of sediment measured by the tool. In older sediments, plate motion may have moved the site latitude since deposition. In this case the remanent anomaly should be interpreted in terms of the paleoinclination rather than the present inclination at the site. When the depositional beds are tilted or the borehole deviates from vertical, the geometrical situation must be established so that the relation between the remanent magnetization and the field in the borehole can be derived. 2.4.2. Is the GHMT polarity affected by overprints? A problem commonly encountered in cores taken by the ODP is the magnetic overprint imparted by the drill string. This overprint is commonly directed vertically downward and radially outward from the centre of core (Roberts et al., 1996; Fuller et al., 1998). The drill string overprint is usually removed by AF demagnetization in peak fields of 20 mT, but sometimes some overprint remains (e.g. Hole 987E, Channell et al., 1999),
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and occasionally the overprint has been strong enough to entirely obscure the remanent magnetization (e.g. Leg 154, Curry et al., 1995). Therefore the question arises: to what extent does the overprint affect the sediment around the borehole measured by the GHMT? The presence of both positive and negative remanent anomaly values that form valid magnetostratigraphies is strong evidence that the original sediment magnetization is not overprinted. A small overprint would be seen as an amplitude reduction in the stable remanent anomaly. The remanent magnetization of the sediment (Jr ) around the borehole can be calculated from the remanent anomaly using Eq. (3) (above), and is illustrated for Hole 1096B in Fig. 5. Whereas the initial core magnetization is overprinted in the downward direction, the in situ magnetization derived from the GHMT logs shows both normal and reversed polarity. The core magnetization after the 20-mT demagnetization step is of lower magnitude than the in situ magnetization, but the two curves have generally the same shape. We interpret these observations to mean that the drill string overprint replaces the low-coercivity remanent magnetization in the cored sediment and usually does not significantly affect the higher-coercivity components. The GHMT data show that the drill string overprint does not extend strongly into the sediment surrounding the borehole, leaving the in situ magnetization largely intact. It is probable that some small volume immediately adjacent to the drill pipe becomes overprinted, but that the larger volume measured by the GHMT retains its original magnetization. 2.4.3. Intensity of the remanent magnetization The varying of intensity of the remanent magnetization within a polarity chron affects the amplitude of the remanent anomaly but not its polarity, therefore the GHMT magnetostratigraphy is affected only to the extent that the greater the remanent magnetization, the more clearly the remanent anomaly will stand out from the noise. Conversely, there can be useful information about the intensity of remanent magnetization in the remanent anomaly, as discussed in Sections 4 and 5. 2.4.4. Viscous remanent magnetization Viscous remanent magnetization (VRM) is the timedependant component of the magnetization, acquired over time in the direction of the ambient magnetic field (Dunlop, 1973). The sediments measured by the GHMT have been sitting in the Brunhes normal polarity magnetic field since 0.78 Ma, and thus will have acquired a VRM in this direction. Because it is unidirectional, much of the VRM is compensated by the baseline offset
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2.4.5. Depth accuracy of the magnetic field and magnetic susceptibility measurements The average upward velocity of the GHMT as it takes measurements is about 600 m/h, and, although the ship heave is compensated to a certain extent, there are oscillations around the average velocity that can result in dm-scale depth discrepancies between the actual and apparent depth of measurement. The susceptibility sensor and the magnetometer are separated by about 5 m, therefore the magnetic field of a particular bed is measured approximately 30 s after the susceptibility, with a different depth discrepancy. In moderately calm seas with normal ship heaves, depth discrepancy is generally smaller than the vertical resolution of the tool, and therefore not a problem. Where the discrepancy is large, there can be some inaccuracy in the remanent anomaly because it is calculated by the subtraction of the susceptibility-derived induced anomaly from the total field. However, it presents a greater problem for the correlation analysis method when the features being correlated are on the scale of the depth discrepancy. Comparison of results from two logging passes provides a check of the effects of depth accuracy.
Fig. 5. A comparison of the in-situ formation NRM with the initial core NRM and the core NRM after AF demagnetization in peak fields of 20 mT. Note that the overprint present in the initial core NRM is not present in the in-situ formation.
2.4.6. Invalid data spikes Brief abrupt shifts to lower magnetic field values (spikes) are observed in the data from approximately 70% of the ODP holes in which the GHMT was run (Table 1). They are attributable to a tool malfunction that occurs when the tool is struck sharply, and are accompanied by drops in the tool voltage log. Therefore, when the voltage drops below a certain value, and invalid magnetic data can be identified and removed from the log (see example in Fig. 5 of Williams et al., 2002). This results in gaps in the remanent anomaly log, but is again a greater problem for the correlation analysis method, for which continuous remanent and induced anomaly logs are necessary. Interpolation is a partial solution in this case. 3. Examples of magnetostratigraphies from ODP GHMT logs
to zero (Section 2.3). However, it is also dependant on the concentration of magnetic minerals, so it will vary in a similar way to the induced magnetization; this is not corrected for in the current scheme. VRM does not have a significant effect on the polarity interpretation of the example data in Fig. 6, though it is possible that it is responsible for the marginally higher normal polarity remanent anomaly amplitudes compared with reversed polarity amplitudes.
Perhaps the most convincing evidence that overprints and other uncertainties generally do not significantly affect the GHMT data comes from examples of successful GHMT magnetostratigraphies (Fig. 6). The classical example of GHMT magnetostratigraphy comes from Hole 884E, in the North Pacific (Fig. 6A, Dubuisson et al., 1995), drilled in clayey diatom ooze, where every polarity subchron from 1.6 to 8.5 Ma in the Cande and Kent (1995) timescale is clearly repre-
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Fig. 6. Examples of remanent anomalies (re-calculated for this review) and their magnetostratigraphic interpretation from ODP Holes 884E (Leg 145, North Pacific), 987E (Leg 162, North Atlantic), and 1084A (Leg 175, eastern South Atlantic). C + K 1995: the Cande and Kent (1995) geomagnetic polarity timescale (GPTS).
sented in the GHMT magnetostratigraphy. The recovered sediments at this site have magnetizations averaging 30 mA/m after 15 mT peak field AF demagnetization, and a drill string overprint was removed by 5–10 mT AF demagnetization. Hole 987E, drilled in strongly magnetized sediment (initial magnetization 10–70 mA/m), gives remanent anomalies in the ±50 nT range (Fig. 6B). The magnetostratigraphy includes the base of the Jaramillo, and the Olduvai and Reunion subchrons. Shipboard passthrough magnetometer data after demagnetization in peak fields of 25 mT is noisy, as a result of remaining drill-string overprint and drilling-related core deformation, and the GHMT data provide a clearer definition of the magnetostratigraphy in this hole (Channell et al., 1999; Higgins et al., 1999). This hole also provides an example of successful correlation analysis. Hole 1084A, drilled in weakly magnetized clay and diatom-rich nannofossil oozes (Wefer et al., 1998), gives remanent anomalies to ±10 nT (Fig. 6C). The magnetization of recovered sediment ranges from 0.1 to 5 mA/m after 20 mT peak field AF demagnetization; a significant drill string overprint remained in the XCB cores after AF demagnetization. The site is located at 25.5◦ S, where the time average magnetic inclination is ∼29◦ , close to the value of 35◦ where the remanent anomaly cannot be observed by the total field measurement in the borehole
(Fig. 2). This, together with the weaker sediment magnetization, provides more of a challenge for the GHMT, but a useful magnetostratigraphy is nevertheless obtained. Holes 998B and 1001A provide an example of successful GHMT magnetostratigraphy by correlation analysis from relatively weakly magnetized sediments, with initial magnetization of 0.2–6 mA/m (see Louvel and Galbrun, 2000). For Hole 998B, a strong overprint prevented a magnetostratigraphy from being obtained from cores, however the GHMT magnetostratigraphy was able to be produced from 20 to 27 Ma which was consistent with biostratigraphy. The Hole 1001A GHMT magnetostratigraphy, from 53 to 69 Ma, was consistent with both the core paleomagnetic results and the biostratigraphy. GHMT magnetostratigraphies are more robust when they are obtained from high latitudes and from finegrained sediments rich in magnetic minerals. This is exemplified by the GHMT logs from Leg 178, offshore of the Antarctic Peninsula, where the polarity transitions can be seen even in the raw total magnetic field log (Williams et al., 2002). GHMT magnetostratigraphies could not be obtained for all holes for a variety of reasons. For example: the Amazon Fan sediments drilled on Leg 155 were entirely Brunhes in age; the GHMT was plagued by invalid data spikes on Leg 184 to the South China Sea; and the induced magnetization dominated the
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remanent magnetization (i.e. a low Koenigsberger ratio) in the carbonate-rich sediments of the Blake Nose on Leg 171B. 4. Can paleointensity be derived from the magnetic logs? The magnitude of the remanent anomaly is proportional to the geomagnetic paleointensity and the concentration of magnetic minerals in the sediment. It also depends upon the type, size, and shape of those minerals and on their host sediment (see Section 5). Thibal
et al. (1995) discuss the plausibility of deriving relative geomagnetic paleointensity from GHMT data by normalizing the remanent anomaly by magnetic susceptibility, and present a favorable comparison of the GHMT-derived relative paleointensity from Hole 884E with the relative paleointensity record of Valet and Meynadier (1994) for 2.7–4.7 Ma. The method is directly analogous to normalizing core paleomagnetic intensity by susceptibility, with susceptibility providing a measure of the concentration of magnetic minerals. All of the tests and assumptions that apply for cores (Tauxe, 1993; Constable and Tauxe, 1996) therefore also apply
Fig. 7. Plots of the Koenigsberger ratio, representing relative paleointensity, from two passes of the GHMT in Hole 1095C. The linear best fit to the data in each of the polarity intervals is also plotted—the paleointensity decays with time in the 10 out of 17 polarity intervals, suggesting a preference for the saw-tooth pattern from 5 to 10 Ma.
T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273
for downhole measurements. It is important to consider that deriving relative paleointensity from the GHMT is more difficult than deriving magnetostratigraphy: magnetostratigaphy requires only that the polarity of the remanent anomaly is correctly identified, whereas for paleointensity, the amplitude of the remanent anomaly must be reliably determined. Thibal et al. also point out that the GHMT relative paleointensity record from Hole 884E follows the sawtooth pattern (where paleointensity decreases from the start to the end of each polarity interval) described by Valet and Meynadier, 1993, for the last 4 Ma. The origin of the saw-tooth pattern has been debated: the principal ideas are that it is geomagnetic in origin (Meynadier et al., 1994), or that it is non-geomagnetic in origin, resulting from delayed magnetization from the subsequent polarity (Mazaud, 1996) or by cumulative viscous remanence (Kok and Tauxe, 1996). Here we use the ratio between the remanent and induced magnetization (the Koenigsberger ratio) as an indicator of relative paleointensity to check if the sawtooth pattern persists back in time to 10 Ma. The GHMT data from Hole 1095B are known to carry a clear polarity signal (Williams et al., 2002), and moreover, parallel records from two separate runs of the GHMT in the hole provide a test of repeatability (Fig. 7). Over this time interval, to my knowledge there are no published paleointensity records to directly compare with. A linear best-fit line has been calculated for 17 polarity intervals from 5 to 10 Ma, with the result that the paleointensity in 10 polarity chrons decays with time (a saw-tooth pattern), 3 increase with time, and 4 are undetermined (are different in the two passes). Thus, there appears to be a preference for the saw-tooth pattern of relative paleointensity variation back to 10 Ma. The relative paleointensities derived from the GHMT should be regarded as ‘quick and dirty’ estimates that, while potentially providing useful and continuous records, lack the back-up checks available for core measurements, and include uncertainties specific to the downhole GHMT measurement.
271
ratio provides a guide to which sediments are likely to give a reliable GHMT-based magnetostratigraphy, and, more generally, a guide to how much lithological variation can be tolerated when constructing paleointensity records. I selected log intervals that had a relatively homogenous sediment type, covering at least 1 Myr, for this analysis (Williams et al., 2002). The ratios have been corrected for latitude by assuming an axial dipole field. The value of the Koenigsberger ratio depends on two main factors: the type of sediment and the paleointensity of the Earth’s magnetic field. The dependence on sediment type is caused in large part by the grain size: larger (multi-domain) magnetic grains have a smaller magnetization (per volume) than smaller (pseudo-single domain) grains, and are less well able to overcome mechanical resistance from the surrounding matrix and align along the magnetic field. Hence, a sediment containing predominantly large ferrimagnetic grains will have a smaller remanence that one containing smaller grains. Lithic clasts that contain magnetic grains are even more unlikely to be able to align along the magnetic field. Because the susceptibility, and hence the induced magnetization, is fairly constant with magnetite grain size (Heider et al., 1996), the Koenigsberger ratio should also be smaller for coarser grained sediments. The ratios presented here vary by about a factor of 10, depending on the sediment type (Fig. 8), with the finer-grained, clay-bearing sediments being the most efficient paleomagnetic recorders. Ratios from just the fine grained sediment types are more constant, varying by a factor of about 2. Paleointensity records are usually generated from this kind of fine grained sediment, and this study offers background information about the variability of the paleomagnetic recording efficiency; it is
5. Efficiency of the depositional remanence acquisition process determined from GHMT data The paleomagnetic recording efficiency of different sediment types can be assessed using the time-averaged Koenigsberger ratio (the ratio of the mean absolute value for the remanent anomaly and the mean value of the induced anomaly) (Parkinson and Barnes, 1985). Such analysis is more difficult on cored sediment because of overprints and demagnetization. The Koenigsberger
Fig. 8. Comparison of time-averaged Koenigsberger ratios in different sediment types (see text).
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important that the recording efficiency is as constant as possible in the particular sediment interval under study. More determinations of the Koenigsberger ratio from different sites will be useful to add weight to these initial analyses, and can provide information for selecting sites which are likely to provide a GHMT-based reversal stratigraphy. Of the sites in Fig. 8, all those with ratios greater than 1 provide GHMT-based magnetostratigraphies, while those with ratios below 1 are marginal or do not provide magnetostratigraphies. 6. Conclusions Examination of the GHMT data obtained by ODP leads to the conclusion that the GHMT can provide valid magnetostratigraphies under a wide range of conditions, including in sediments with only moderate remanent magnetizations (down to approximately 1 mA/m). The principal problem with the ODP GHMT data has been the presence of invalid data spikes in the total magnetic field record; however, these data are easily identified and removed from the record and usually need not prevent a magnetostratigraphy from being obtained. Unlike recovered core samples, the downhole measurement is not significantly affected by drilling overprint, enabling a magnetostratigraphy to be obtained even when the paleomagnetic record from cores is compromised. Additionally, the GHMT provides information about the paleomagnetic recording efficiency of different types of sediment, something that is more difficult to obtain from recovered sediment due to demagnetization and the overprint. The GHMT magnetic field and susceptibility data also provide the means to obtain ‘quick and dirty’ relative paleointensity records in sediments that are well magnetized, enabling identification of sections suitable for further study on recovered sediment. In summary, the GHMT has been a valuable tool as part of the suite of downhole measurements made in the ODP, and its data has been an integrated part of the paleomagnetic studies providing magnetostratigraphies to aid in dating the sediment column. Acknowledgements This research used data provided by the Ocean Drilling Program (ODP). ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. The logging scientists and Schlumberger engineers who collected the GHMT data onboard the JOIDES Resolution are gratefully acknowledged, as is Veronique Louvel, who was responsible for
the GHMT data after it was collected. Thoughtful and constructive reviews by Richard Jarrard and Yoshifumi Nogi helped improve the manuscript.
References Bosum, W., Scott, J.H., 1988. Interpretation of magnetic logs in basalt Hole 418A. In: Salisbury, M.H., Scott, J.H. (Eds.), Proc. ODP, Sci. Results, vol. 102. College Station, TX (Ocean Drilling Program), pp. 77–95. Cande, S.C., Kent, D.V., 1995. Revised calibration of the geomagnetic polarity timescale for the Late Cretaceous and Cenozoic. J. Geophys. Res. 100, 6093–6095. Channell, J.E.T., Smelror, M., Jansen, E., Higgins, S.M., Lehman, B., Eidvin, T., Solheim, A., 1999. Age models for glacial fan deposits off East Greenland and Svalbard (Sites 986 and 987). In: Proc. ODP, Sci. Results, vol. 162, pp. 149–166. Constable, C., Tauxe, L., 1996. Towards absolute calibration of sedimentary paleointensity records. Earth Planet. Sci. Lett. 143, 269–274. Gallet, Y., Courtillot, V., 1989. Modeling magnetostraigraphy in a borehole. Geophysics 54 (8), 973–983. Curry, W.B., Shackleton, N.J., Richter, C., et al., 1995. Proc. ODP Init. Repts. 154, College Station, TX (Ocean Drilling Program). Bouisset, P.M., Augustin, A.M., 1993. Borehole magnetostratigraphy, absolute age dating, and correlation of sedimentary rocks, with examples from the Paris Basin, France. AAPG Bulletin, vol. 77, pp. 569–587. Dubuisson, G., Thibal, J., Barthes, V., Pocachard, J., Pozzi, J.-P., 1995. Downhole magnetic logging in sediments during Leg 145: usefulness and magnetostratigraphic interpretation of the logs at Site 884. In: Proc. ODP, Sci. Results, vol. 145, pp. 455–468. Dunlop, D.J., 1973. Theory of the magnetic viscosity of lunar and terrestrial rocks. Rev. Geophys. Space Phys. 11 (4), 855–901. Fuller, M., Hastedt, M., Herr, B., 1998. Coring-induced magnetization of recovered sediment. In: Proc. ODP, Sci. Results, vol. 145, College Station, TX (Ocean Drilling Program), pp. 455– 468. Hamano, Y., Kinoshita, H., 1990. Magnetization of the oceanic crust inferred from magnetic logging in Hole 395A. In: Detrick, R.S., Honnorez, J., Bryan, W.B., Juteau, T. (Eds.), Proc. ODP, Sci. Results, vols. 106/109. College Station, TX (Ocean Drilling Program), pp. 223–229. Heider, F., Zitzelsberger, A., Fabian, K., 1996. Magnetic susceptibility and remanent coercive force in grown magnetite crystals from 0 1 m to 6 mm. Phys. Earth Planet. Interiors 93, 239– 256. Higgins, S.M., Kreitz, S., King, T., Goldberg, D., 1999. Data report: Magnetic polarity and susceptibility measurements from the geological high-resolution magnetometer tool at Sites 984, 986 and 987. In: Proc. ODP, Sci. Results, vol. 162, pp. 265–269. Ito, H., Nogi, Y., Larson, R.L., 1995. Magnetic reversal stratigraphy of Jurassic oceanic crust from Hole 801C downhole magnetometer measurements. In: Haggerty, J.A., Premoli-Silva, I., Rack, F.R., McNutt, M.K. (Eds.), Proc. ODP, Sci. Results, vol. 144. College Station, TX (Ocean Drilling Program), pp. 641–647. Kinoshita, H., Furuta, T., Pariso, J.E., 1989. Downhole magnetic field measurements and paleomagnetism, Hole 504B Costa Rica Ridge. In: Becker, K., Sakai, H. (Eds.), Proc. ODP, Sci. Results, vol. 111. College Station, TX (Ocean Drilling Program), pp. 147–156.
T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273 Kok, Y.S., Tauxe, L., 1996. Saw-toothed pattern of sedimentary paleointensity records explained by cumulative viscous remanence. Earth Planet. Sci. Lett. 144, E9–E14. Louvel, V., Galbrun, B., 2000. Magnetic polarity sequences from downhole measurements in ODP Holes 998B and 1001A, Leg 165, Caribbean Sea. Mar. Geophys. Res. 21, 561–577. Mazaud, A., 1996. Sawtooth variation in magnetic intensity profiles and delayed acquisition of magnetization in deep sea cores. Earth Planet. Sci. Lett. 139, 379–386. Meynadier, L., Valet, J.-P., Bassinot, F., Shackleton, N.J., 1994. Assymetrical saw-tooth pattern of the geomagnetic field intensity from equatorial sediments in the Pacific and Indian Oceans. Earth Planet. Sci. Lett. 126, 109–127. Nogi, Y., Tarduno, J.A., Sager, W.W., 1995. Magnetization of seamount-derived sediments from Site 869 inferred from downhole magnetometer logs. In: Winterer, E.L., Sager, W.W., Firth, J., Sinton, J.M. (Eds.), Proc. ODP, Sci. Results, vol. 143. College Station, TX (Ocean Drilling Program), pp. 373–379. Pariso, J.E., Scott, J.H., Kikawa, E., Johnson, H.P., 1991. A magnetic logging study of Hole 735B gabbros at the Southwest Indian Ridge. In: Von Herzen, R.P., Robinson, P.T. (Eds.), Proc. ODP, Sci. Results, vol. 118. College Station, TX (Ocean Drilling Program), pp. 309–322. Parker, R.L., Daniell, G.J., 1979. Interpretation of borehole magnetometer data. J. Geophys. Res. 10, 5467–5479. Parkinson, W.D., Barnes, C.D., 1985. In situ determination of Koenigsberger ratio. Aust. J. Earth Sci. 32, 1–5. Ponomarev, V.N., Nechoroshkov, V.L., 1984. Downhole magnetic measurements in oceanic crustal Hole 395A on the Mid-Atlantic Ridge. Init. Repts. DSDP, 78B. U.S. Govt. Printing Office, Washington. Ponomarev, V.N., and Nechoroshkov, V.L., 1983. First measurements of the magnetic field within the ocean crust; Deep Sea Drilling Project Legs 68 and 69. Init. Repts. DSDP 69. U.S. Govt. Printing Office, Washington.
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Pozzi, J.-P., Martin, J.P., Pocachard, J., Feinberg, H., Galdeano, A., 1988. In situ magnetostratigraphy: interpretation of magnetic logging in sediments. Earth Planet. Sci. Lett. 88, 357–373. Pozzi, J.-P., Barthes, V., Thibal, J., Pocachard, J., Lim, M., Thomas, T., Pages, G., 1993. Downhole magnetostratigraphy in sediments: Comparison with the paleomagnetism of a core. J. Geophys. Res. 98, 7939–7957. Roberts, A.P., Stoner, J.S., Richter, C., 1996. Coring-induced magnetic overprints and limitations of the long-core paleomagnetic measurement technique: some observations from Leg 160, eastern Mediterranean Sea. In: Proc. ODP Init. Repts. 160, pp. 497– 505. Roperch, P., Barth`es, V., Pocachard, J., Collot, J.-Y., Chabernaud, T., 1994. Magnetic logging and in-situ magnetostratigraphy: a field test. In: Proc. ODP, Sci. Results, vol. 134, pp. 577–589. Tarduno, J.A., Duncan, R.A., Scholl, D.W., et al., 2002. Proc. ODP Init. Repts. 197, College Station TX 77845-9547, USA (Available from: Ocean Drilling Program, Texas A&M University (CD-ROM)). Tauxe, L., 1993. Sedimentary records of relative paleointensity of the geomagnetic field; theory and practice. Rev. Geophys. 31, 319–354. Thibal, J., Pozzi, J.-P., Barth´es, V., Dubuisson, G., 1995. Continuous record of geomagnetic field intensity between 4.7 and 2.7 Ma from downhole measurements. Earth Planet. Sci. Lett. 136, 541– 550. Valet, J.-P., Meynadier, L., 1993. Geomagnetic field intensity and reversals during the past 4 million years. Nature 366, 91–95. Vibert-Charbonnel, P., 1996. Methodes de traitment des mesures magnetiques en forage pour la datation haute-resolution des series sedimentaires. Thesis. L’Ins. Nat Polytech. de Grenoble. Wefer, G., Berger, W.H., Richter, C., et al., 1998. Proc. ODP Init. Repts., vol. 175, College Station, TX (Ocean Drilling Program). Williams, T., Louvel, V., Lauer-Leredde, C., 2002. Magnetic polarity stratigraphy from downhole logs, West Antarctic Peninsula, ODP Leg 178. In: Proc. ODP, Sci. Results, vol. 178, pp. 1–23.
Magnetostratigraphy from downhole measurements in ODP holes Trevor Williams ∗ Borehole Research Group, Lamont-Doherty Earth Observatory, 61 Route 9W, Palisades, NY 10964, United States Received 24 May 2004; received in revised form 14 July 2005; accepted 22 September 2005
Abstract The “Geological High-resolution Magnetic Tool” (GHMT) was used on 17 ODP Legs to obtain downhole logs of magnetic field strength and magnetic susceptibility. The remanent magnetization of the sediment around the borehole is derived from these measurements, and used to produce magnetostratigraphies. This review of the ODP GHMT data indicates that reliable magnetostratigraphies can usually be obtained from reasonably well-magnetized (above approximately 10−3 A/m) fine-grained sediments. Additionally, the GHMT data indicates that a drill-string overprint, common in data collected on cores, does not significantly affect the sediments surrounding the borehole. This allows us to use the remanent and induced magnetization data to obtain estimates of relative paleointensity, and when averaged over time in homogenous sediment, the paleomagnetic recording efficiency of that sediment. © 2006 Elsevier B.V. All rights reserved. Keywords: Magnetostratigraphy; ODP; Downhole measurements; Paleomagnetism; Paleointensity
1. Introduction The “Geological High-resolution Magnetic Tool” (GHMT) was developed as a cooperative venture between Commissariat a l’Energie Atomique–Laboratoire d’Electronique, de Technologie et d’Instrumentation (CEA–LETI), the Ecole Normale Superieur, and the Total Oil Company in France, and has been operated in the Ocean Drilling Program (ODP) by Schlumberger starting in 1990 (ODP Leg 134). The motivation for developing the tool was to provide magnetostratigraphic dating of formations that had only limited biostratigraphic dating from either spot coring or cuttings. Magnetostratigraphic dating requires good time-sampling
∗
Tel.: +1 845 365 8626. E-mail address: [email protected].
0031-9201/$ – see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.pepi.2005.09.016
of the stratigraphic section, and the GHMT provides this by coverage of the whole open hole interval. In conventional paleomagnetic measurements on cores, Uchannels, and samples, several checks on the quality of the data are available, such as sequential demagnetization and visual inspection of the sediment for deformation or inhomogeneities. Additionally, paleomagnetic measurements are made in low-field environments, so there is no induced component to the sample’s magnetization. In measurements made downhole, demagnetization and visual checks are not possible, and the sediment has induced and viscous magnetizations in addition to the remanent magnetization which carries the magnetostratigraphy. Despite the differences, these concerns can be assessed: questions about the GHMT data and its interpretation are addressed in this review. Compilation of the ODP GHMT data set and comparison to paleomagnetism measured on cores recovered from the
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same holes indicates that valid magnetostratigraphies are obtained by the GHMT in reasonably well-magnetized sediment sections, for example down to approximately 10−3 A/m. Given a strong magnetization, the GHMT can provide information about relative paleointensity and the paleomagnetic recording efficiency of different sediment types. The GHMT measures the magnetic susceptibility in the formation surrounding the borehole using induction, and measures the total magnetic field in the borehole using a proton-precession magnetometer. In order to access the magnetostratigraphy, the remanent field anomaly must be isolated from the total field measured in the borehole. The measured field is dominated by the Earth’s dipole field, but is also contributed to by local fields due to the drill pipe, basement rocks, inhomogeneities in the formation, and the induced magnetization of the sediment. All of these must be subtracted to leave the remanent field anomaly: the quality of the magnetostratigraphy depends on how large the remanent anomaly is compared to the errors and noise in the estimations of the other fields. The theory describing the magnetic field in a borehole caused by the magnetization of the rocks surrounding the borehole was described by Parker and Daniell (1979) to enable interpretation of future downhole magnetometer observations in terms of the heterogeneity of magnetization in basaltic ocean crust. Pozzi et al. (1988) and Gallet and Courtillot (1989) developed the theory for horizontal and dipping magnetized layers with the aim of deriving the magnetostratigraphy of sedimentary successions. Initial field tests of the GHMT are described for the Permian Lod`eve Basin in the south of France (Pozzi et al., 1988), and for the Paris Basin (Bouisset and Augustin, 1993; Pozzi et al., 1993). GHMT results for ODP holes are presented for Leg 134 (Roperch et al., 1994), Leg 145 (Dubuisson et al., 1995), Leg 162 (Higgins et al., 1999), Leg 165 (Louvel and Galbrun, 2000), and Leg 178 (Williams et al., 2002). The GHMT was also run on ODP Legs 155, 160, 171B, 175, 184, 188, and 189, as summarized in Table 1. GHMT magnetic susceptibility records were obtained on ODP Legs 154, 167, 177, 181, and 182. Besides the GHMT proton-precession magnetometer, a variety of downhole fluxgate magnetometers have been used in ODP and DSDP boreholes for tool orientation or investigation of the magnetization of igneous ocean crust. Schlumberger’s general purpose inclinometry tool (GPIT) is a three-axis fluxgate tool that has been used on most ODP legs since Leg 126 mainly for orientation of downhole resistivity images, but also occasionally for paleomagnetism, for example in oceanic
basement in Hole 504B (Kinoshita et al., 1989) and in Hole 801C (Ito et al., 1995). A Russian three-axis fluxgate magnetometer was first used on Legs 68 and 69, and successfully detected reversals between basalt flows in Hole 395A on DSDP Leg 78B (Ponomarev and Nechoroshkov, 1983, 1984). A Japanese three-axis fluxgate magnetometer was also run in Hole 395A on Leg 109 (Hamano and Kinoshita, 1990). The Japanese magnetometer was also run on Leg 143 in both the sediments of Hole 869 (Nogi et al., 1995a) and the basalts of Holes 865 and 866 (Nogi et al., 1995b). A gyroscopeoriented three-axis fluxgate magnetometer, developed in Hannover, Germany, was run in Hole 418A, Leg 102, yielding magnetic characteristics of basalts and paleomagnetic pole positions (Bosum and Scott, 1988). Pariso et al. (1991), compared results from two downhole fluxgate magnetometers, one from the University of Washington and one from the US Geological Survey, in the Gabbros of Hole 735B. More recently, the Gottingen Borehole Magnetometer was run in alternating basalts and volcaniclastic sediments in Hole 1203A of Leg 197 (Tarduno et al., 2002). It comprises a three-axis fluxgate magnetometer together with an optical angular rate sensor, which, in theory, is accurate enough to determine the orientation of the tool independently from the magnetic reference frame, enabling the horizontal direction of the field to be determined. The advantage of the proton precession magnetometer over the fluxgate tools is its accuracy (0.1 nT), relative insensitivity to temperature variations, and its lack of measurement drift. The three-axis fluxgate magnetometers provide the vertical and horizontal components of the field, but the absence of directional information in the total field measurement of the GHMT is not generally a barrier to polarity determination, as discussed below. The GHMT magnetic susceptibility logs provide a valuable lithological indicator sensitive to paleoceanographic and paleoclimatic change, as well as a good parameter for correlation between cores and downhole logs. This paper is concerned solely with the paleomagnetic applications of the GHMT, and the paleoclimatic applications are not discussed further here. 2. Derivation of polarity of remanent magnetization from the total magnetic field measured in a borehole 2.1. Remanent magnetization The remanent magnetization (Jr ) of sediments is typically acquired during deposition by preferential alignment of the magnetic grains along the magnetic field
Table 1 Magnetic field and magnetic susceptibility logs were obtained by the GHMT in the holes listed in this table Leg
Hole
Location
Depth range (m)
Age
Induced anomaly range (nT)
1990 1992 1992
134 145 145
831B 883F 884E
Bougainville Guyot Detroit Seamount Detroit Seamount
100–700 220–720 90–680
Oligocene–Pleistocene Oligocene–l. Pliocene l. Miocene–l. Pliocene
30 70
1994 1994 1995 1995 1995
155 155 160 160 162
931B 933A 966F 967E 986C
Amazon Fan Amazon Fan Eratosthenes Seamount Eratosthenes Seamount Svalbard Margin
90–240 90–210 100–340 100–580 100–360
l. Pleistocene l. Pleistocene Eocene–l. Pliocene l. Pliocene Pleistocene
1995
162
987E
Greenland Margin
100–300
1996 1996 1997 1997 1997 1997 1997 1997 1998
165 165 171B 171B 171B 175 175 175 178
998B 1001A 1050C 1051A 1052E 1081A 1082A 1084A 1095B
Cayman Rise Nicaraguan Rise Blake Nose Blake Nose Blake Nose Walvis Ridge Walvis Basin Cape Basin Antarctic Peninsula
1998
178
1096C
1998 1999 1999 1999 2000 2000 2000 2000
178 184 184 184 188 189 189 189
1103A 1144A 1146A 1148A 1166A 1168A 1170D 1172D
Remanent anomaly range (nT)
Invalid data spikes
Reference
30 100
None None Present
Roperch et al. (1994) Dubuisson et al. (1995) Dubuisson et al. (1995), this paper
25 20 3 3 80
20 20 2 8 100
None Present None Present Present
l. Pliocene–Pleistocene
100
150
Present
210–430 200–400 125–585 140–590 240–680 100–400 90–550 80–590 110–550
Oligocene–e. Miocene Cretaceous–Paleocene Cretaceous–Eocene Cretaceous–Eocene Cretaceous–Paleocene l. Miocene–l. Pliocene l. Pliocene–e. Pleistocene l. Pliocene–l. Pleistocene l. Miocene–e. Pliocene
5 6 12 10 10 8 8 8 100
8 10 4 3 3 8 12 25 200
None None Present Present Present Present Present Present Present
Antarctic Peninsula
360–500
e. Pliocene
100
200
None
Antarctic Peninsula N Margin S China Sea N Margin S China Sea N Margin S China Sea Prydz Bay West Tasmania Slope South Tasman Rise East Tasman Plateau
100–240 120–440 250–585 400–705 40–380 140–720 550–780 210–760
l. Pliocene–l. Pleistocene l. Pleistocene m. Miocene–e. Pliocene Oligocene–e. Miocene Eocene-Pliocene e. Oligocene–l. Miocene Eocene Cretaceous–l. Miocene
350 6 6 5 300 35 35 50
100 6 6 2 50 10 10 10
Abundant Abundant Abundant Abundant Present Present None Present
Higgins et al. (1999), Channell et al. (1999) Higgins et al. (1999), Channell et al. (1999) Louvel and Galbrun (2000) Louvel and Galbrun (2000)
This paper This paper, Williams et al. (2002) This paper, Williams et al. (2002)
T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273
Year
The amplitude ranges of the remanent and induced anomalies are approximate values for decameter scale variation. The number of invalid data spikes are classified here as none (if the spikes are absent or negligible), present (if spikes are present in all or part of the log and can be removed to leave good data), or abundant (if the spikes dominate the log and prevent analysis). GHMT-based magnetostratigraphies can generally be produced from holes where the remanent anomaly is equal to or higher than the induced anomaly. The GHMT data is available from http://iodp.ldeo.columbia.edu/DATA/index.html. 263
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direction. The remanent magnetization contributes to the field anomaly in the borehole (Bfr ), according to the following equations (Pozzi et al., 1988) (Fig. 1A): Bfrx =
µ 0
2
· Jrx
(1)
Bfrz = −µ0 · Jrz
(2)
where Bfrx and Bfrz are the components of Bfr in the x (∼north) and z (down) directions. The Earth’s main field, Br , is much larger than the field anomaly due to the remanent magnetization (in the example in Fig. 1, 44,000 nT compared to 114 nT), so the total field, B0 , and the Earth’s field, Br , are sub-parallel (the angle between the fields is 0.104◦ , Fig. 1B). Therefore, to a very close approximation, it is the projection of the remanent anomaly Bfrp along B0 that contributes to the measured B0 . This is the basis for determining polarity using the intensity of the total field rather than its directional change: whereas an ±80 nT change in total field is easy to measure, the directional change in the field is very difficult to measure, especially from a tool that is free to rotate and is subject to impacts in the borehole. Bfrp is calculated by (Pozzi et al., 1988): Bfrp =
µ0 · J r 2
· (1 − 3 sin2 I)
(3)
Fig. 2. Dependence on the field inclination of the remanent anomaly, Bfrp , in the borehole due to a 0.1 A/m remanent magnetization. Note that at an inclination of ±35◦ , no remanent anomaly is observed, because Bfr is perpendicular to the earth’s field (see Fig. 1B).
where I is the inclination of the remanent magnetization. Note that at an inclination of 35◦ , the field due to the remanent magnetization Bfrp is perpendicular to the total field, and the anomaly cannot be observed with total field measurement (Fig. 2). 2.2. Induced magnetization The sediment also carries an induced magnetization (Ji ) along the direction of the present-day field. Ji can be calculated from the magnetic induction H and the magnetic susceptibility κ (measured by the GHMT) (Eqs. (4)
Fig. 1. An example showing the remanent and induced magnetizations (Jr and Ji ) in the formation and the remanent and induced field anomalies they cause in the borehole (Bfr and Bfi ). Given a sediment magnetization, Jr , of 0.1 A/m, the observed anomaly in the borehole, Bfr , is 80 nT (see text). The example is from the southern hemisphere, at the latitude of Hole 1096B, where the field inclination is −60.6◦ . Note that the fields in the borehole do not have the same direction as the sediment magnetizations that produce them.
T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273
and (5)). The field anomaly in the borehole Bfip caused by the induced magnetization Ji can then be calculated by using the equivalent of Eq. (3). B0 Ji = ·κ (4) µ0 B 0 = µ0 · H
(5)
where: B is the magnetic induction, unit T; H is the magnetic field, unit A/m; J is the magnetization, volume normalized, unit A/m; κ is the magnetic susceptibility, volume normalized, unitless (×10−6 SI units); µ0 is the permeability of nonmagnetic material = 4π × 10−7 (in SI units). 2.3. Deriving the GHMT polarity from the measured total field The measured total field is the vector sum of a number of different contributing fields, and the remanent anomaly is found by progressively subtracting the other individual fields from the total (see Fig. 3). B0 (z, t) = Br (z) + Ba (z) + Bfr (z) + Bfi (z) + Bt (z, t) (6) All the components of B0 are a function of depth, z; the time-dependant magnetic field, Bt , generated in the ionosphere, also contributes. Br , the field generated in the Earth’s core, forms the largest component of B0 by about two orders of magnitude. It is estimated from the International Geomagnetic Reference Field (IGRF). Ba is the regional magnetic field: sea-floor and other crustal anomalies. The field caused by the metal (magnetic) drill pipe (Bpipe ) is also included within Ba ; it decays with distance (r) away from the pipe, Bpipe = a/r3 . The pipe effect is determined by iterative forward modeling until a satisfactory fit to the data is obtained. Bt is the time dependant field, originating in the magnetosphere and ionosphere, which varies on timescales of minutes to days. Two passes of the GHMT are usually run so that the repeatability can be checked; for the most part the differences are slight and do not significantly affect the data. However, negative ‘spikes’ to low field values occur in many of the GHMT logs, and these spikes do not repeat from pass to pass: they originate in the tool itself and are not a valid recording of the field. They are accompanied by a drop in the tool voltage, and the invalid magnetic field data may be removed from the log on the basis of the voltage (Williams et al., 2002). Bfip , the induced field anomaly, is calculated using Eqs. (3)–(5), with values for B0 and I taken from the IGRF values for the site.
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Even after Br , Ba , and Bfip have been subtracted from the measured total field B0 to leave Bfrp , the baseline value of the remanent field anomaly, Bfrp can be offset from zero. This offset can be caused by diurnal variation, local discrepancies from the IGRF, proximity to basement crustal anomalies, etc. However, we know a priori that over enough normal and reverse polarity intervals Bfrp will average out to near zero. Therefore, we can correct for this baseline offset by subtracting a constant value or a linear fit, resulting in a remanent anomaly Bfrp that is positive in normal polarity sediment and negative in reversed polarity sediment. Thus, the first method for determining polarity is the sign of the remanent anomaly, which is often sufficiently good to provide a basis for magnetostratigraphy. A second method, correlation analysis, overcomes the problem of a variable baseline offset (Pozzi et al., 1993; Vibert-Charbonnel, 1996; Louvel and Galbrun, 2000). The principle employed by the correlation analysis method is illustrated in Fig. 4. Both Jr and Ji are dependant on the concentration of ferrimagnetic minerals in the sediment, and this concentration varies. Jr and Ji correlate in normal polarity intervals, and the linear regression line has a positive gradient, whereas in reversed polarity intervals Jr is inversely correlated with Ji and the linear regression line has a negative gradient. The same applies to Bfrp and Bfip , on which the correlation analyses are actually performed. The linear regressions are applied to successive depth intervals (windows) of various thickness (1.5, 2.5, 4.5, 8, and 13 m windows are presented in Fig. 3). Linear regressions with a correlation coefficient below 0.5 are not plotted. Prior to the correlation analyses, Bfrp and Bfip are smoothed with an 11 sample (1.5 m) Hanning filter, so that the Bfrp and Bfip logs have comparable vertical resolution. Results of the two methods are plotted side-by-side in Fig. 3: the sign of the remanent anomaly gives a more continuous polarity record, while correlation analysis gives a better reading close to the drill pipe. 2.4. Constraints on GHMT polarity determination In order to evaluate the qualitative confidence that should be applied to GHMT magnetostratigraphies, the assumptions involved in the analysis are discussed below, followed by presentation of examples of demonstrably successful GHMT magnetostratigraphies in ODP. The majority of the issues below affect only the amplitude of the remanent anomaly and not its polarity, therefore leaving the GHMT magnetostratigraphy minimally affected.
266 T. Williams / Physics of the Earth and Planetary Interiors 156 (2006) 261–273 Fig. 3. An example from Hole 1096B showing the derivation of the remanent anomaly from the measured magnetic susceptibility and total field logs. The original measured logs are susceptibility and the total field. The induced anomaly is calculated from the susceptibility, and then the remanent anomaly is calculated by subtraction of the induced anomaly, the pipe effect, and a constant value from the total field. Polarity can be determined from (a) the sign of the remanent anomaly and from (b) the correlation analysis method (positive gradients in black, negative gradients in green). In each case a polarity column is given, with black indicating normal polarity, green reversed polarity, and white undetermined polarity. The depth window employed in the correlation analyses are (from left to right) 1.5, 2.5, 4.5, 8, and 13 m. Paleomagnetic data from Hole 1096B after demagnetization in peak fields of 20 mT and the GPTS of Cande and Kent (1995) are given for comparison.
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Fig. 4. Remanent (Bfrp ) vs. induced (Bfip ) anomalies for the interval from 400 to 420 mbsf in Hole 1096C, covering the Gilbert–Gauss polarity reversal, illustrating the principles of the correlation analysis method. Correlation between Bfrp and Bfip indicates normal polarity, and anti-correlation indicates reversed polarity.
2.4.1. Inclination of the remanent magnetization Both the amplitude and polarity of the remanent anomaly depend on the inclination of the remanent magnetization (Fig. 2), so the inclination must be known or estimated. In the simplest case of geologically young horizontally bedded sediments, the remanent magnetization will be either parallel or anti-parallel to the Earth’s axial dipole field. Most of the ODP boreholes logged with the GHMT approximate to this case. Paleomagnetic secular variation will generally average out to the axial dipole field over the volume of sediment measured by the tool. In older sediments, plate motion may have moved the site latitude since deposition. In this case the remanent anomaly should be interpreted in terms of the paleoinclination rather than the present inclination at the site. When the depositional beds are tilted or the borehole deviates from vertical, the geometrical situation must be established so that the relation between the remanent magnetization and the field in the borehole can be derived. 2.4.2. Is the GHMT polarity affected by overprints? A problem commonly encountered in cores taken by the ODP is the magnetic overprint imparted by the drill string. This overprint is commonly directed vertically downward and radially outward from the centre of core (Roberts et al., 1996; Fuller et al., 1998). The drill string overprint is usually removed by AF demagnetization in peak fields of 20 mT, but sometimes some overprint remains (e.g. Hole 987E, Channell et al., 1999),
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and occasionally the overprint has been strong enough to entirely obscure the remanent magnetization (e.g. Leg 154, Curry et al., 1995). Therefore the question arises: to what extent does the overprint affect the sediment around the borehole measured by the GHMT? The presence of both positive and negative remanent anomaly values that form valid magnetostratigraphies is strong evidence that the original sediment magnetization is not overprinted. A small overprint would be seen as an amplitude reduction in the stable remanent anomaly. The remanent magnetization of the sediment (Jr ) around the borehole can be calculated from the remanent anomaly using Eq. (3) (above), and is illustrated for Hole 1096B in Fig. 5. Whereas the initial core magnetization is overprinted in the downward direction, the in situ magnetization derived from the GHMT logs shows both normal and reversed polarity. The core magnetization after the 20-mT demagnetization step is of lower magnitude than the in situ magnetization, but the two curves have generally the same shape. We interpret these observations to mean that the drill string overprint replaces the low-coercivity remanent magnetization in the cored sediment and usually does not significantly affect the higher-coercivity components. The GHMT data show that the drill string overprint does not extend strongly into the sediment surrounding the borehole, leaving the in situ magnetization largely intact. It is probable that some small volume immediately adjacent to the drill pipe becomes overprinted, but that the larger volume measured by the GHMT retains its original magnetization. 2.4.3. Intensity of the remanent magnetization The varying of intensity of the remanent magnetization within a polarity chron affects the amplitude of the remanent anomaly but not its polarity, therefore the GHMT magnetostratigraphy is affected only to the extent that the greater the remanent magnetization, the more clearly the remanent anomaly will stand out from the noise. Conversely, there can be useful information about the intensity of remanent magnetization in the remanent anomaly, as discussed in Sections 4 and 5. 2.4.4. Viscous remanent magnetization Viscous remanent magnetization (VRM) is the timedependant component of the magnetization, acquired over time in the direction of the ambient magnetic field (Dunlop, 1973). The sediments measured by the GHMT have been sitting in the Brunhes normal polarity magnetic field since 0.78 Ma, and thus will have acquired a VRM in this direction. Because it is unidirectional, much of the VRM is compensated by the baseline offset
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2.4.5. Depth accuracy of the magnetic field and magnetic susceptibility measurements The average upward velocity of the GHMT as it takes measurements is about 600 m/h, and, although the ship heave is compensated to a certain extent, there are oscillations around the average velocity that can result in dm-scale depth discrepancies between the actual and apparent depth of measurement. The susceptibility sensor and the magnetometer are separated by about 5 m, therefore the magnetic field of a particular bed is measured approximately 30 s after the susceptibility, with a different depth discrepancy. In moderately calm seas with normal ship heaves, depth discrepancy is generally smaller than the vertical resolution of the tool, and therefore not a problem. Where the discrepancy is large, there can be some inaccuracy in the remanent anomaly because it is calculated by the subtraction of the susceptibility-derived induced anomaly from the total field. However, it presents a greater problem for the correlation analysis method when the features being correlated are on the scale of the depth discrepancy. Comparison of results from two logging passes provides a check of the effects of depth accuracy.
Fig. 5. A comparison of the in-situ formation NRM with the initial core NRM and the core NRM after AF demagnetization in peak fields of 20 mT. Note that the overprint present in the initial core NRM is not present in the in-situ formation.
2.4.6. Invalid data spikes Brief abrupt shifts to lower magnetic field values (spikes) are observed in the data from approximately 70% of the ODP holes in which the GHMT was run (Table 1). They are attributable to a tool malfunction that occurs when the tool is struck sharply, and are accompanied by drops in the tool voltage log. Therefore, when the voltage drops below a certain value, and invalid magnetic data can be identified and removed from the log (see example in Fig. 5 of Williams et al., 2002). This results in gaps in the remanent anomaly log, but is again a greater problem for the correlation analysis method, for which continuous remanent and induced anomaly logs are necessary. Interpolation is a partial solution in this case. 3. Examples of magnetostratigraphies from ODP GHMT logs
to zero (Section 2.3). However, it is also dependant on the concentration of magnetic minerals, so it will vary in a similar way to the induced magnetization; this is not corrected for in the current scheme. VRM does not have a significant effect on the polarity interpretation of the example data in Fig. 6, though it is possible that it is responsible for the marginally higher normal polarity remanent anomaly amplitudes compared with reversed polarity amplitudes.
Perhaps the most convincing evidence that overprints and other uncertainties generally do not significantly affect the GHMT data comes from examples of successful GHMT magnetostratigraphies (Fig. 6). The classical example of GHMT magnetostratigraphy comes from Hole 884E, in the North Pacific (Fig. 6A, Dubuisson et al., 1995), drilled in clayey diatom ooze, where every polarity subchron from 1.6 to 8.5 Ma in the Cande and Kent (1995) timescale is clearly repre-
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Fig. 6. Examples of remanent anomalies (re-calculated for this review) and their magnetostratigraphic interpretation from ODP Holes 884E (Leg 145, North Pacific), 987E (Leg 162, North Atlantic), and 1084A (Leg 175, eastern South Atlantic). C + K 1995: the Cande and Kent (1995) geomagnetic polarity timescale (GPTS).
sented in the GHMT magnetostratigraphy. The recovered sediments at this site have magnetizations averaging 30 mA/m after 15 mT peak field AF demagnetization, and a drill string overprint was removed by 5–10 mT AF demagnetization. Hole 987E, drilled in strongly magnetized sediment (initial magnetization 10–70 mA/m), gives remanent anomalies in the ±50 nT range (Fig. 6B). The magnetostratigraphy includes the base of the Jaramillo, and the Olduvai and Reunion subchrons. Shipboard passthrough magnetometer data after demagnetization in peak fields of 25 mT is noisy, as a result of remaining drill-string overprint and drilling-related core deformation, and the GHMT data provide a clearer definition of the magnetostratigraphy in this hole (Channell et al., 1999; Higgins et al., 1999). This hole also provides an example of successful correlation analysis. Hole 1084A, drilled in weakly magnetized clay and diatom-rich nannofossil oozes (Wefer et al., 1998), gives remanent anomalies to ±10 nT (Fig. 6C). The magnetization of recovered sediment ranges from 0.1 to 5 mA/m after 20 mT peak field AF demagnetization; a significant drill string overprint remained in the XCB cores after AF demagnetization. The site is located at 25.5◦ S, where the time average magnetic inclination is ∼29◦ , close to the value of 35◦ where the remanent anomaly cannot be observed by the total field measurement in the borehole
(Fig. 2). This, together with the weaker sediment magnetization, provides more of a challenge for the GHMT, but a useful magnetostratigraphy is nevertheless obtained. Holes 998B and 1001A provide an example of successful GHMT magnetostratigraphy by correlation analysis from relatively weakly magnetized sediments, with initial magnetization of 0.2–6 mA/m (see Louvel and Galbrun, 2000). For Hole 998B, a strong overprint prevented a magnetostratigraphy from being obtained from cores, however the GHMT magnetostratigraphy was able to be produced from 20 to 27 Ma which was consistent with biostratigraphy. The Hole 1001A GHMT magnetostratigraphy, from 53 to 69 Ma, was consistent with both the core paleomagnetic results and the biostratigraphy. GHMT magnetostratigraphies are more robust when they are obtained from high latitudes and from finegrained sediments rich in magnetic minerals. This is exemplified by the GHMT logs from Leg 178, offshore of the Antarctic Peninsula, where the polarity transitions can be seen even in the raw total magnetic field log (Williams et al., 2002). GHMT magnetostratigraphies could not be obtained for all holes for a variety of reasons. For example: the Amazon Fan sediments drilled on Leg 155 were entirely Brunhes in age; the GHMT was plagued by invalid data spikes on Leg 184 to the South China Sea; and the induced magnetization dominated the
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remanent magnetization (i.e. a low Koenigsberger ratio) in the carbonate-rich sediments of the Blake Nose on Leg 171B. 4. Can paleointensity be derived from the magnetic logs? The magnitude of the remanent anomaly is proportional to the geomagnetic paleointensity and the concentration of magnetic minerals in the sediment. It also depends upon the type, size, and shape of those minerals and on their host sediment (see Section 5). Thibal
et al. (1995) discuss the plausibility of deriving relative geomagnetic paleointensity from GHMT data by normalizing the remanent anomaly by magnetic susceptibility, and present a favorable comparison of the GHMT-derived relative paleointensity from Hole 884E with the relative paleointensity record of Valet and Meynadier (1994) for 2.7–4.7 Ma. The method is directly analogous to normalizing core paleomagnetic intensity by susceptibility, with susceptibility providing a measure of the concentration of magnetic minerals. All of the tests and assumptions that apply for cores (Tauxe, 1993; Constable and Tauxe, 1996) therefore also apply
Fig. 7. Plots of the Koenigsberger ratio, representing relative paleointensity, from two passes of the GHMT in Hole 1095C. The linear best fit to the data in each of the polarity intervals is also plotted—the paleointensity decays with time in the 10 out of 17 polarity intervals, suggesting a preference for the saw-tooth pattern from 5 to 10 Ma.
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for downhole measurements. It is important to consider that deriving relative paleointensity from the GHMT is more difficult than deriving magnetostratigraphy: magnetostratigaphy requires only that the polarity of the remanent anomaly is correctly identified, whereas for paleointensity, the amplitude of the remanent anomaly must be reliably determined. Thibal et al. also point out that the GHMT relative paleointensity record from Hole 884E follows the sawtooth pattern (where paleointensity decreases from the start to the end of each polarity interval) described by Valet and Meynadier, 1993, for the last 4 Ma. The origin of the saw-tooth pattern has been debated: the principal ideas are that it is geomagnetic in origin (Meynadier et al., 1994), or that it is non-geomagnetic in origin, resulting from delayed magnetization from the subsequent polarity (Mazaud, 1996) or by cumulative viscous remanence (Kok and Tauxe, 1996). Here we use the ratio between the remanent and induced magnetization (the Koenigsberger ratio) as an indicator of relative paleointensity to check if the sawtooth pattern persists back in time to 10 Ma. The GHMT data from Hole 1095B are known to carry a clear polarity signal (Williams et al., 2002), and moreover, parallel records from two separate runs of the GHMT in the hole provide a test of repeatability (Fig. 7). Over this time interval, to my knowledge there are no published paleointensity records to directly compare with. A linear best-fit line has been calculated for 17 polarity intervals from 5 to 10 Ma, with the result that the paleointensity in 10 polarity chrons decays with time (a saw-tooth pattern), 3 increase with time, and 4 are undetermined (are different in the two passes). Thus, there appears to be a preference for the saw-tooth pattern of relative paleointensity variation back to 10 Ma. The relative paleointensities derived from the GHMT should be regarded as ‘quick and dirty’ estimates that, while potentially providing useful and continuous records, lack the back-up checks available for core measurements, and include uncertainties specific to the downhole GHMT measurement.
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ratio provides a guide to which sediments are likely to give a reliable GHMT-based magnetostratigraphy, and, more generally, a guide to how much lithological variation can be tolerated when constructing paleointensity records. I selected log intervals that had a relatively homogenous sediment type, covering at least 1 Myr, for this analysis (Williams et al., 2002). The ratios have been corrected for latitude by assuming an axial dipole field. The value of the Koenigsberger ratio depends on two main factors: the type of sediment and the paleointensity of the Earth’s magnetic field. The dependence on sediment type is caused in large part by the grain size: larger (multi-domain) magnetic grains have a smaller magnetization (per volume) than smaller (pseudo-single domain) grains, and are less well able to overcome mechanical resistance from the surrounding matrix and align along the magnetic field. Hence, a sediment containing predominantly large ferrimagnetic grains will have a smaller remanence that one containing smaller grains. Lithic clasts that contain magnetic grains are even more unlikely to be able to align along the magnetic field. Because the susceptibility, and hence the induced magnetization, is fairly constant with magnetite grain size (Heider et al., 1996), the Koenigsberger ratio should also be smaller for coarser grained sediments. The ratios presented here vary by about a factor of 10, depending on the sediment type (Fig. 8), with the finer-grained, clay-bearing sediments being the most efficient paleomagnetic recorders. Ratios from just the fine grained sediment types are more constant, varying by a factor of about 2. Paleointensity records are usually generated from this kind of fine grained sediment, and this study offers background information about the variability of the paleomagnetic recording efficiency; it is
5. Efficiency of the depositional remanence acquisition process determined from GHMT data The paleomagnetic recording efficiency of different sediment types can be assessed using the time-averaged Koenigsberger ratio (the ratio of the mean absolute value for the remanent anomaly and the mean value of the induced anomaly) (Parkinson and Barnes, 1985). Such analysis is more difficult on cored sediment because of overprints and demagnetization. The Koenigsberger
Fig. 8. Comparison of time-averaged Koenigsberger ratios in different sediment types (see text).
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important that the recording efficiency is as constant as possible in the particular sediment interval under study. More determinations of the Koenigsberger ratio from different sites will be useful to add weight to these initial analyses, and can provide information for selecting sites which are likely to provide a GHMT-based reversal stratigraphy. Of the sites in Fig. 8, all those with ratios greater than 1 provide GHMT-based magnetostratigraphies, while those with ratios below 1 are marginal or do not provide magnetostratigraphies. 6. Conclusions Examination of the GHMT data obtained by ODP leads to the conclusion that the GHMT can provide valid magnetostratigraphies under a wide range of conditions, including in sediments with only moderate remanent magnetizations (down to approximately 1 mA/m). The principal problem with the ODP GHMT data has been the presence of invalid data spikes in the total magnetic field record; however, these data are easily identified and removed from the record and usually need not prevent a magnetostratigraphy from being obtained. Unlike recovered core samples, the downhole measurement is not significantly affected by drilling overprint, enabling a magnetostratigraphy to be obtained even when the paleomagnetic record from cores is compromised. Additionally, the GHMT provides information about the paleomagnetic recording efficiency of different types of sediment, something that is more difficult to obtain from recovered sediment due to demagnetization and the overprint. The GHMT magnetic field and susceptibility data also provide the means to obtain ‘quick and dirty’ relative paleointensity records in sediments that are well magnetized, enabling identification of sections suitable for further study on recovered sediment. In summary, the GHMT has been a valuable tool as part of the suite of downhole measurements made in the ODP, and its data has been an integrated part of the paleomagnetic studies providing magnetostratigraphies to aid in dating the sediment column. Acknowledgements This research used data provided by the Ocean Drilling Program (ODP). ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. The logging scientists and Schlumberger engineers who collected the GHMT data onboard the JOIDES Resolution are gratefully acknowledged, as is Veronique Louvel, who was responsible for
the GHMT data after it was collected. Thoughtful and constructive reviews by Richard Jarrard and Yoshifumi Nogi helped improve the manuscript.
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