Tectonophysics 418 (2006) 145 – 162
www.elsevier.com/locate/tecto
Magnetic fabric variations in Mesozoic black shales, Northern
Siberia, Russia: Possible paleomagnetic implications
Martin Chadima a,⁎, Petr Pruner a , Stanislav Šlechta a , Tomáš Grygar b , Ann M. Hirt c
a
Institute of Geology, Academy of Sciences, Rozvojová 135, CZ-16502 Prague, Czech Republic
b
Institute of Inorganic Chemistry, Academy of Sciences, CZ-25068 Rež, Czech Republic
c
Institut für Geophysik, ETH-Hönggerberg, CH-8093 Zürich, Switzerland
Received 13 June 2005; received in revised form 20 October 2005; accepted 5 December 2005
Available online 28 February 2006
Abstract
A 28-m-long section situated on the coast of the Arctic Ocean, Russia (74°N, 113°E) was extensively sampled primarily for the
purpose of magnetostratigraphic investigations across the Jurassic/Cretaceous boundary. The section consists predominantly of
marine black shales with abundant siderite concretions and several distinct siderite cemented layers. Low-field magnetic
susceptibility (k) ranges from 8 × 10− 5 to 2 × 10− 3 SI and is predominantly controlled by the paramagnetic minerals, i.e. ironbearing chlorites, micas, and siderite. The siderite-bearing samples possess the highest magnetic susceptibility, usually one order of
magnitude higher than the neighboring rock. The intensity of the natural remanent magnetization (M0) varies between 1 × 10− 5 and
6 × 10− 3 A/m. Several samples possessing extremely high values of M0 were found. There is no apparent correlation between the
high k and high M0 values; on the contrary, the samples with relatively high M0 values possess average magnetic susceptibility and
vice versa. According to the low-field anisotropy of magnetic susceptibility (AMS), three different groups of samples can be
distinguished. In the siderite-bearing samples (i), an inverse magnetic fabric is observed, i.e., the maximum and minimum principal
susceptibility directions are interchanged and the magnetic fabric has a distinctly prolate shape. Triaxial-fabric samples (ii),
showing an intermediate magnetic fabric, are always characterized by high M0 values. It seems probable that the magnetic fabric is
controlled by the preferred orientation of paramagnetic phyllosilicates, e.g., chlorite and mica, and by some ferromagnetic mineral
with anomalous orientation in relation to the bedding plane. Oblate-fabric samples (iii) are characterized by a bedding-controlled
magnetic fabric, and by moderate magnetic susceptibility and M0 values. The magnetic fabric is controlled by the preferred
orientation of phyllosilicate minerals and, to a minor extent, by a ferrimagnetic fraction, most probably detrital magnetite.
Considering the magnetic fabric together with paleomagnetic component analyses, the siderite-bearing, and the high-NRM samples
(about 15% of samples) were excluded from further magnetostratigraphic research.
© 2006 Elsevier B.V. All rights reserved.
Keywords: AMS; AARM; High-field anisotropy; Inverse magnetic fabric; Siderite; Jurassic/Cretaceous boundary
1. Introduction
Recent magnetostratigraphic investigations across
the Jurassic/Cretaceous (J/K) boundary carried out on
⁎ Corresponding author. Tel.: +42 272 690 115.
E-mail address: chadima@sci.muni.cz (M. Chadima).
0040-1951/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2005.12.018
three representative sections in the Tethyan realm of
Europe yielded comparable results (Houša et al., 1999,
2004). In order to establish a magnetostratigraphic
correlation with the Boreal realm, a section across the
J/K boundary was studied in the Eastern Russian
Arctic. As a complementary technique to magnetostratigraphy, the measurements of the anisotropy of the
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M. Chadima et al. / Tectonophysics 418 (2006) 145–162
low-field magnetic susceptibility (AMS) were performed throughout the section. Unexpectedly, several
different orientations and shapes of magnetic fabric
ellipsoid were found. The AMS of any rock is
dependent on the intrinsic magnetic susceptibility,
volume fraction, and degree of preferred orientation
of the individual rock-constituent minerals. For this
reason the AMS reflects not only the differences in the
rock fabric but also any mineralogical variations within
the studied rock units (e.g. Robion et al., 1995). Such
variations, if caused by the presence of different types
of ferromagnetic minerals (e.g. Lehman et al., 1996),
may lead to erroneous magnetostratigraphic interpretations and should be, therefore, thoroughly understood.
In sedimentary rocks unaffected by tectonic ductile
deformation, the so-called ‘normal magnetic fabric’ is
usually observed. The normal magnetic fabric is
characterized by magnetic foliation oriented parallel
to the bedding, and magnetic lineation being roughly
parallel to the near-bottom water current direction or,
in special cases, perpendicular to it (Hamilton and
Rees, 1970). The magnitude of the sedimentary
magnetic fabric is relatively low and the fabric is
distinctly oblate. Despite the predominant occurrence
of normal magnetic fabrics, ‘inverse magnetic fabrics’
are observed in the sedimentary rocks containing
minerals with an inverse relationship between magnetic axes and shape and/or crystallographic axes (e.g.
Rochette, 1988). In rocks possessing the inverse
fabric, the principal magnetic axes are inverted, thus,
both magnetic foliation and magnetic lineation are
perpendicular to the bedding plane and the fabric is
distinctly prolate (Rochette et al., 1992). Inverse
magnetic fabric can be carried by single-domain
magnetite (Potter and Stephenson, 1988; Rochette et
al., 1999; Borradaile and Gauthier, 2001), tourmaline
(Rochette et al., 1994; Ferré and Améglio, 2000),
cordierite, goethite (Lehman et al., 1996), or by different iron-bearing carbonates (Ellwood et al., 1986;
Rochette, 1988; Ellwood et al., 1989; Ihmlé et al., 1989;
Winkler et al., 1996; Hounslow, 2001; de Wall and
Warr, 2004).
The purpose of the present paper is to interpret the
different types of magnetic fabric found in the studied
sedimentary rocks. An integrated magnetic fabric
approach (AMS, high-field anisotropy, anisotropy of
anhysteretic remanent magnetization) is combined with
standard magnetic and non-magnetic mineral identification techniques. The understanding of the mineralogical control on magnetic fabric is used to help
discriminate which samples are unsuitable for a reliable
magnetostratigraphy.
2. Geological setting and sampling
The magnetostratigraphic and magnetic fabric investigation was carried out in a sedimentary complex
exposed on the coast of the Laptev Sea (Zakharov et al.,
1983). The sedimentary sequence was deposited in a
foreland basin on the passive continental margin
bounding the Siberian Craton to the north. The rocks
from the foreland basin were subsequently deformed by
folding and thrusting in the Early Cretaceous (Drachev
et al., 1998).
The studied outcrop is located on the Nordvik
Peninsula, Urdyuk-Khaya Cape, west coast of the
Anabar Bay (73°54′N, 113°04′E, Fig. 1a). The rocks
representing the supposed J/K boundary interval are
exposed as a broad syncline with the dip of strata
changing progressively from subhorizontal to approximately 60°. Several faults with adjacent zones of brittle
deformation cross-cut the outcrop. The section is
composed of marine mudstones and siltstones, which
we will denote as shales, with abundant concretions and
several distinct cemented layers (Fig. 1b). A newly
formed siderite was observed macroscopically in the
concretions and cemented layers. In addition to siderite,
several horizons with small pyrite nodules and thin
pyrite layers were present within the shales (Fig. 1b).
The main constituents of the studied shales, which have
been identified by microscopic observations, are rock
fragments, quartz, feldspars, micas, chlorites, glauconite
and pyrite (Zakharov et al., 1983).
Overall 370 oriented samples were taken from three
individual sections (Fig. 1b). The H- and D-sections
were sampled above and below the supposed J/K
boundary, respectively. This boundary is marked by a
4- to 6-cm-thick phosphate limestone horizon with
high contents of iridium and other noble metals
(Zakharov et al., 1993). Due to the occurrence of the
fault zones, a continuation of the D-section was
sampled several hundred meters away, denoted as Msection. Approximately 2-m overlap interval between
the D- and M-section was sampled in order to ensure
the correct linking (Fig. 1b). Sample names include
the acronym of the respective section plus figure
expressing the distance of the sample from the
supposed J/K boundary in cm. In the central part,
dense sampling at 2–4 cm spacing was made (Dsection), whereas the marginal parts were sampled at
approximately 10-cm intervals (H- and M-sections).
The samples were cut as 2 × 2 × 2 cm cubes directly
from the outcrop, using a diamond-coated wheel saw.
In order to prevent further disintegration of the very
loose sediment, each sample was pressed into a plastic
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
147
Fig. 1. Location (a) and the lithological log (b) of the studied section on the Nordvik Peninsula, Northern Siberia, Russia. The log is labeled in cm,
with negative or positive numbers denoting positions above or below the supposed Jurassic/Cretaceous boundary, respectively. Lithological log
drawn by M. Mazuch.
cube. A significant limitation to using plastic cases is
that samples cannot be thermally demagnetized or
subjected to any thermal rock magnetic analysis.
3. Laboratory techniques and data processing
3.1. Identification of magnetic minerals
Temperature variations of magnetic susceptibility
from room temperature up to 700 °C were measured
with an Agico KLY-4S Kappabridge coupled with a
temperature control apparatus CS-3. All the measurements were performed on powdered samples in air
atmosphere, with a heating rate of approximately 10 °C/
min. The curves of the temperature dependence of
magnetic susceptibility allow the resolution of the room
temperature susceptibility into its paramagnetic and
ferromagnetic components based on the hyperbolafitting method (Hrouda, 1994; Hrouda et al., 1997).
In order to test for the presence of pyrrhotite, bulk
susceptibility was measured as a function of applied AC
field with an Agico KLY-4S Kappabridge. The field
intensity was gradually increased from 2 to 450 A/m.
An isothermal remanent magnetization (IRM) was
acquired progressively using a Magnetic Measurements
MMPM10 pulse magnetizer in fields up to 1 T. The
acquired IRM was measured after each imparted field
with an Agico JR-5 spinner magnetometer. Backfield
demagnetization was done using a pulse magnetizer.
The S-ratio was calculated according to Bloemendal et
al. (1992):
S−0:3
T
¼ ½−ðIRM−0:3T =SIRMÞ þ 1=2
where SIRM is the saturation IRM, which was imparted
applying an IRM in one direction in a maximum field of
1 T, and IRM− 0.3T is the magnetization after the
application of 300 mT field in the opposite direction.
All the rock magnetic measurements were carried out in
the Paleomagnetic laboratory, Institute of Geology,
Prague, Czech Republic.
The X-ray diffraction (XRD) was done using a
Siemens D5005 diffractometer in the Institute of
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Inorganic Chemistry, Řež, Czech Republic. The diffractometer is equipped with a Cu X-ray source with a
wavelength λ = 1.54 Å. The diffraction scans were made
on powder samples in the range of 2θ angle from 2° to
80° with angular velocity of 1°/min.
Backscattered electron (BSE) images on polished
sections were done using a Cameca SX-100 Electron
Probe MicroAnalyzer having four wave-dispersive
spectrometers (WDS) in the Institute of Geology,
Prague.
3.2. Natural remanent magnetization
The natural remanent magnetization (NRM) was
investigated to study the magnetic polarity for magnetostratigraphic purposes in the Geological Research
Center, Potsdam, Germany. Stepwise alternating field
(AF) demagnetization up to a maximum field of 100 mT
was performed with a 2G Enterprises degausser system.
After each demagnetization step, the remanent magnetization was measured on a 2G Enterprises cryogenic
magnetometer. Characteristic remanent magnetization
(ChRM) directions were identified from vector-component diagrams using principal component analysis
(Kirschvink, 1980). The mean ChRM directions were
analyzed using the statistics on sphere (Fisher, 1953).
3.3. Magnetic fabric
The AMS and the anisotropy of anhysteretic
remanent magnetization (AARM) was studied in the
Paleomagnetic laboratory in Prague. The AMS was
measured with an Agico KLY-4S Kappabridge with an
alternating field intensity of 300 A/m and operating
frequency of 875 Hz. The AARM was measured using
an Agico AF demagnetizer/magnetizer LDA-3/AMU-1
and an Agico JR-6 spinner magnetometer. The remanent
magnetization was imparted in six pairs of antiparallel
directions with a DC field of 100 μT and AC field of
50 mT. Using this measurement design, the ‘hard’
magnetization components, which cannot be demagnetized, are reasonably eliminated (Jelínek, 1993). A
tumbling system was used to demagnetize the samples
after each pair of magnetizing positions.
The high-field magnetic anisotropy (HFA) was
measured using a high-field torque magnetometer
(Bergmüller et al., 1994) in the Laboratory for Natural
Magnetism, Institute of Geophysics, ETH, Zürich,
Switzerland. In the high-field torque magnetometer, a
sample is measured along three mutually perpendicular
axes with an angular increment of 20° in four different
fields. These fields were strong enough to saturate
ferrimagnetic minerals and allowed the separation of
paramagnetic and ferrimagnetic contribution to the
magnetic anisotropy using the method of MartínHernández and Hirt (2001).
The magnetic anisotropy (either AMS, AARM, or
HFA) can be mathematically described as a symmetric
second rank tensor which can be visualized as an
ellipsoid. Its semi-axes lengths, k1 ≥ k2 ≥ k3, are termed
principal values and their orientations, K1, K2, K3, are
denoted as principal directions. In the framework of
magnetic fabric, the maximum direction (K1) defines a
magnetic lineation while the plane perpendicular to
minimum direction (K3) and containing maximum and
intermediate directions (K1, K2) defines a magnetic
foliation.
Quantitatively, the magnetic anisotropy is presented
in terms of the mean susceptibility, k (applies to the
AMS), the degree of anisotropy P (AMS, AARM), the
shape parameter, T (AMS, AARM), and the difference
shape factor, U (AMS, HFA) defined as follows:
k = (k1 + k2 + k3) / 3
P = k1 / k3
T = 2ln(k2 / k3) / ln(k1 / k3) − 1
U = (2k2 − k1 − k3) / (k1 − k3) =
[2(k2 − k) − (k1 − k) − (k3 − k)] / [(k1 − k) − (k3 − k)]
(Nagata, 1961)
(Nagata, 1961)
(Jelínek, 1981)
(Jelínek, 1981)
where (k1 − k) ≥ (k2 − k) ≥ (k3 − k) are the deviatoric principal values.
The P parameter indicates the magnitude of the
magnetic anisotropy and the T parameter indicates the
shape of the anisotropy ellipsoids; it varies from − 1
(perfectly prolate magnetic fabric) to +1 (perfectly
oblate magnetic fabric). The employment of the
infrequently used difference shape factor, U, arises
from the intended comparison of the AMS with the
HFA. As the torque magnetometer measures only the
deviatoric component of anisotropy, only the U
parameter based on the susceptibility differences can
be calculated. The shape difference factor, U, is
analogous to the T shape factor, ranging from − 1 to 1;
negative or positive values indicate prolate- or oblatefabrics, respectively.
4. Magnetic mineralogy
A number of samples have been subjected to
magnetic and non-magnetic mineral identification
techniques, and three types of magnetic behavior can
be identified. For this reason, data from three representative samples are presented. The basic parameters of
NRM and AMS for these samples are summarized in
Table 1.
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Table 1
The basic parameters of NRM and AMS for representative samples
Sample
D0653
D0907
D1245
MDF
M0
−2
5.01 × 10
3.31 × 10− 4
1.74 × 10− 4
25.6
18.7
26.5
k
−6
201.6 × 10
206.0 × 10− 6
1055.3 × 10− 6
P
T
Inc K1
Inc K2
Inc K3
1.092
1.055
1.150
0.022
0.995
− 0.990
8.4
13.3
81.7
68.7
6.4
1.9
19.4
75.2
8.1
M0, intensity of natural remanent magnetization in A/m; MDF, medium destructive field in mT; k, mean susceptibility in 10− 6 SI; P, magnitude of
AMS; T, shape parameter of AMS; Inc K1, Inc K2, Inc K3, inclinations of maximum, intermediate, and minimum susceptibility directions,
respectively, in bedding coordinate system (tilt correction).
4.1. Temperature variation of magnetic susceptibility
During the laboratory heating of the powder samples,
the magnetic susceptibility decreases as a function of
increasing temperature following a paramagnetic hyperbola according to the Curie–Weiss law (Fig. 2).
Depending on the sample, the hyperbolic decrease is
maintained up to the temperature of 240 or 300 °C.
Susceptibility resolution in the low-temperature hyperbolic interval showed that room temperature bulk
magnetic susceptibility is dominated by the paramagnetic fraction. Paramagnetic contribution to the room
temperature bulk susceptibility ranges from approximately 80% to 100% (Fig. 2).
Two different sample behaviors are observed above
the temperature of 240 or 300 °C. For some samples, the
susceptibility starts to increase after 240 °C, first slowly,
and above 290 °C, rapidly. The slope of the curve
changes again in the temperature interval between 340
and 390 °C where a gentle increase is evident (Fig. 2,
D0653). Above 390 °C, the magnetic susceptibility
reaches values an order of magnitude higher than
original room temperature magnetic susceptibility.
Above 500 °C, the magnetic susceptibility starts to
decrease rapidly, and later above 580 °C, follows a
hyperbolic curve.
For the second group of samples, the shoulder in the
susceptibility increase in the 240–390 °C interval is not
present. Magnetic susceptibility starts to increase
rapidly at temperatures above 300 °C (Fig. 2, D0907,
D1245) reaching maximum values of one or two orders
of magnitude higher than the original room temperature
magnetic susceptibility. The susceptibility starts to
decrease above the temperature of 520–545 °C, and
follows again a hyperbolic course in the interval above
approximately 580 °C (D0907) or 625 °C (D1245).
4.2. X-ray diffraction and electron microscopy
The diffraction scans vary according to the studied
sample. The pronounced peaks corresponding to the Xray diffractions on various structural planes of quartz,
feldspars (albite and microcline), muscovite, chlorite
(clinochlore), siderite, pyrite, gypsum, and kaolinite
were identified (Fig. 3). The semi-quantitative phase
composition estimates based on the relative intensity
Fig. 2. Temperature variations of bulk magnetic susceptibility. Test samples were heated in the air atmosphere with a heating rate of 10 °C/min. The
resolution of the room susceptibility into the paramagnetic (p) and ferromagnetic ( f ) component with the standard deviations (in parenthesis) is
presented for each curve.
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Fig. 3. The X-ray powder diffraction scans for three test samples in the range of 2θ angle 2°–80° with angular velocity 1°/min. The main diffraction
peaks of the rock-constituent minerals are labeled: quartz (q), muskovite (m), clinochlore (c), siderite (s), pyrite (p), gypsum (g), albite (a), and
kaolinite (k).
ratios referring to corundum standard are presented in
Table 2. In the BSE images the zonal nodules, corroded
euhedral crystals and framboides are observed in
phyllosilicate matrix (Fig. 4). The atomic weight ratios
derived from WDS analyses of nodules, crystals and
framboides are close to the composition of pyrite.
magnetic susceptibility remains constant during gradual
increase of the AC field from 2 to 450 A/m (Fig. 5b).
5. Magnetic fabric and NRM
5.1. Theoretical models of inverse and anomalous
magnetic fabric
4.3. Ferromagnetic mineralogy
The origin of the inverse, and intermediate magnetic
fabrics was reviewed and modeled by Rochette et al.
(1992, 1999), and Ferré (2002). In these models, the
terms normal, inverse, intermediate, and anomalous
magnetic fabric were defined according to the relationship of the principal anisotropy directions to the
principal directions of structural features. ‘Normal’ is
used to describe a magnetic fabric whose K1 direction is
parallel to a structural lineation (either flow, current,
stretching direction, or apparent lineation), and whose
K3 direction is normal to a structural foliation (e.g.,
plane of flattening, flow or bedding plane, Fig. 6, N).
Strictly following the directional definition, the normal
Several methods of rock magnetic analysis were used
to identify the ferromagnetic mineralogy of the rocks.
The acquisition of IRM showed that all samples reached
magnetic saturation by 200 to 500 mT (Fig. 5a).
Backfield demagnetization of the saturation IRM
(SIRM) revealed a coercivity of remanence, Hcr, between 30 to 50 mT in general. One sample had a Hcr
around 53 mT (Fig. 5a, D1245).
In order to test for the presence of pyrrhotite, bulk
susceptibility was measured as a function of applied AC
field (Worm et al., 1993, Pokorný et al., 2004). No field
dependence is observed in the test samples and bulk
Table 2
The semi-quantitative phase composition estimates (%) of the rock-constituent minerals on the basis of the relative intensity XRD ratios referring to
corundum standard
Sample
Quartz
Pyrite
Albite
Muscovite
Clinoclore
D0653
D0907
D1245
27.6
25.5
43.4
18.2
2.0
4.9
3.0
2.4
7.2
37.1
44.5
21.0
8.1
8.8
7.5
Siderite
16.1
Gypsum
Kaolinite
Microcline
5.9
9.4
5.6
1.9
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
151
Fig. 4. Backscattered electron images of the polished sections of the representative pyrite-bearing sample with different types of pyrite (p) in
phyllosilicate matrix. (a) Pyrite nodules, corroded euhedral crystals and framboides, (b) large corroded pyrite crystal.
magnetic fabric can be in principle oblate, triaxial, or
prolate in symmetry. In the ‘inverse’ magnetic fabric, the
maximum and minimum anisotropy axes are inverted
relative to the normal fabric (Fig. 6, I). The origin of
intermediate fabric has been modeled by mixing the
normal and inverse fabric end-members (Rochette et al.,
1992; Ferré, 2002). Depending on the anisotropy degree
and the symmetry of the end-members, different types
of intermediate fabrics can theoretically develop. No
intermediate fabric develops when a normal and inverse
end-member with the inverted symmetry are mixed.
When the oblate-normal and prolate-inverse endmembers are mixed the ‘false normal’ (K1 and K2 are
interchanged, Fig. 6, FN), and later, the ‘intermediate’
(further interchange of K2 and K3, Fig. 6, IM1) fabrics
develop. When the prolate-normal and prolate-inverse
end-members are mixed the ‘intermediate’ (K2 ↔ K3,
Fig. 6, IM2) and the ‘false inverse’ (further exchange of
K1 and K2, Fig. 6, FI) fabrics develop. Anomalous
fabric is characterized by a random distribution of
magnetic axes (Fig. 6, A).
When no structural lineation is observed in the
field, the orientation of magnetic fabric can be solely
related to the structural foliation, e.g. bedding plane in
undeformed sedimentary rocks, using the inclination of
principal anisotropy directions. In such a simplified
scheme, the ‘false normal’, and ‘false inverse’ fabrics
(cf. Rochette et al., 1999) cannot be distinguished. In
order to fully evaluate all possible orientations of
magnetic fabrics, the ternary diagram interrelating
inclinations of all three principal anisotropy directions
in bedding coordinates is introduced (Fig. 6). In such a
ternary diagram, the normal, and inverse magnetic
fabrics plot in the close vicinity of the lower left and
Fig. 5. (a) An isothermal remanent magnetization (IRM) for the three test samples acquired progressively in fields up to 1 T and backfield IRM
acquired in the field up to 1 T in opposite direction. (b) The bulk magnetic susceptibility measured as a function of applied AC field with the intensity
gradually increased from 2 to 450 A/m.
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Fig. 6. A model ternary diagram interrelating inclinations of the principal anisotropy directions (K1, K2, K3) in bedding coordinate system. The 30°
and 60° threshold angles subdivide the diagram into nine fields. Schematic stereoplots depicting the orientation of principal directions are drawn for
each field. Different types of magnetic fabrics as defined by Rochette et al. (1992, 1999), and Ferré (2002) are presented outside of diagram. Square,
triangle, and circle symbols represent maximum, intermediate, and minimum anisotropy directions, respectively.
lower right vertices, respectively, whereas the intermediate fabrics plot close to the upper vertex. Samples
plotted in the hexagonal area in the center of the
diagram can be regarded as anomalous. The advantage
of the presented diagram is that the different transitions
between the end-member magnetic fabrics (normal,
inverse, and intermediate) can be quantitatively
described.
5.2. Low-field magnetic anisotropy
Magnetic fabric (AMS) of the studied shales is
more or less coaxial with the bedding. For the majority
of samples, the magnetic lineation (K1) is subparallel
to the bedding plane or, in some cases, perpendicular
to it. No distinct preferred orientation of magnetic
lineation within the bedding plane is observed (Fig.
7a). The absence of the cluster-like distribution of
magnetic lineations suggests that magnetic lineation
reflects neither paleocurrent direction nor the incipient
stage of deformation when magnetic lineations tend to
cluster in the direction perpendicular to the beddingparallel shortening (Kligfield et al., 1983) or parallel to
stretching direction in extensional basins (Sagnotti et
al., 1994; Mattei et al., 1997; Cifelli et al., 2005). The
direction of K3 (pole to magnetic foliation) is usually
sub-normal to the bedding plane or, in some cases,
subparallel to it or randomly oriented (Fig. 7b).
Although the low-field magnetic susceptibility varies
widely, ranging from 84 × 10− 6 to 2030 × 10− 6 SI, the
susceptibility values are lower than 250 × 10− 6 SI for the
majority of samples (Fig. 7c). The low susceptibility
values may imply that the paramagnetic minerals
dominate the magnetic susceptibility.
The shape of the AMS ellipsoid was plotted as a
function of the anisotropy degree (Fig. 7d). Three
distinct groups of samples can be distinguished: i)
samples possessing prolate anisotropy (T b − 0.6, henceforth denoted as the prolate-fabric samples) and a wide
range of anisotropy degrees (1 b P b 1.16), ii) samples
possessing triaxial shape (− 0.5 b T b 0.6, the triaxialfabric samples) with less variable anisotropy degree
(generally 1.03 b P b 1.09), and iii) samples characterized by the oblate shape (T N 0.6, the oblate-fabric
samples) and by wide range of anisotropy degrees
(1 b P b 1.16). The above-described division is also
reflected by the magnetic susceptibility (Fig. 7d); the
highest susceptibility (usually, k N 300 × 10− 6 SI) is
observed in the group of the prolate-fabric samples,
whereas the oblate-fabric samples possess significantly
lower susceptibility (usually, k b 300 × 10− 6 SI). The
lowest magnetic susceptibility is possessed by the
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
153
Fig. 7. The anisotropy of low-field magnetic susceptibility (AMS) for the entire set of samples. The maximum (a) and minimum (b) susceptibility
directions are plotted in the equal-area projection on the lower-hemisphere in bedding coordinate system (tilt correction). Frequency distribution of
the susceptibility values (c), relationship between anisotropy degree and shape of anisotropy ellipsoid (P–T plot) and magnetic susceptibility as
grayscales (d), anisotropy degree (P) as a function of magnetic susceptibility (e), and ternary diagram relating the inclination angles of maximum,
intermediate, and minimum susceptibility directions in the bedding coordinate system with the shape parameter (T) of anisotropy ellipsoids as
grayscales (f).
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triaxial-fabric samples (usually, k b 200 × 10− 6 SI). No
apparent correlation between the magnetic susceptibility
and the degree of anisotropy is observed (Fig. 7e).
As already evident from Fig. 7a, b, d, at least two
different types of magnetic fabric (normal and inverse)
can be distinguished. The orientation of the principal
susceptibility axes is in the close relationship to the
shape of the susceptibility ellipsoids (Fig. 7f). The
majority of normal-fabric or inverse-fabric samples
possess oblate or prolate shape, respectively. Despite the
presence of triaxial-fabric samples (Fig. 7d) almost no
pure intermediate fabric (cf., Rochette et al., 1999) is
present but there are some samples that display a gradual
transition from normal to intermediate fabric. The
transitional samples possess both oblate and triaxial
susceptibility ellipsoids (Fig. 7d). The group of inverse
fabric samples is more distinctly delimited (Fig. 7f).
5.3. High-field anisotropy and anisotropy of magnetic
remanence
For the representative oblate-, triaxial-, and prolatefabric samples, the HFA (6 samples) and AARM (18
samples) were measured in order to separate the
paramagnetic and ferromagnetic component of magnetic
anisotropy. All the examined samples show a very
similar linear dependence of the torque force as a
function of the square of the applied field (Fig. 8). The
2Θ term of the Fast Fourier Transformation for all three
perpendicular planes intersects at the origin, therefore
the resolution of the HFA into paramagnetic (henceforth
Fig. 8. Amplitude of the 2Θ-term as a function of the square of the
applied field (B2) for the three perpendicular measurement planes
(squares, triangles, and circles) for sample D1362 where black
symbols represent the coefficients of the cosine term and open
symbols those of the sine term.
denoted as HFP) and ferrimagnetic components shows
that the ferrimagnetic fraction, if present, appears to be
insignificant (Martín-Hernández and Hirt, 2001).
All the samples selected for the AARM measurement
can be magnetized and demagnetized with at least two
orders of magnitude difference in magnetization between the magnetized and demagnetized states. Such a
difference is a prerequisite for the successful application
of the AARM method. The calculated remanent tensors
represent the anisotropy carried by the ferromagnetic
grains characterized by a particular coercivity, size and
shape (Jackson et al., 1988).
The obtained data indicate that magnetic subfabrics
carried by the paramagnetic and ferromagnetic grains
are usually coaxial with the AMS. In some cases, only
one pair of the respective principal directions (maximum
or minimum) is subparallel. For the description of the
orientation AARM or HFP subfabrics, the terms normal,
inverse, intermediate, and anomalous magnetic fabric
will be used in the relationship to the AMS in the
following text. Hereafter, the normal AARM or HFP
fabric is referred to the case when all three respective
principal directions of the AARM or HFP are parallel to
the AMS principal directions. Similarly, the inverse
AARM or HFP fabric is characterized by the interchange of the maximum and minimum principal
directions of the AARM or HFP and the AMS.
In the group of the prolate-fabric samples, the HFP
fabric is normal with respect to the AMS, i.e. all three
principal directions of the HFP and AMS are subparallel
(Fig. 9, samples D1010, D1069). Both the AMS and
HFP fabrics are inverse with respect to the bedding
plane. The AARM is inverse or false inverse to the
AMS, and normal to the bedding plane (Fig. 9).
In the group of the triaxial-fabric samples, the HFP
fabric is false normal (Fig. 9, D0593, D1058), or
anomalous with respect to the AMS (D1362). With
respect to the bedding plane, the HFP fabric is normal
(D1058, D1362) or inverse (D0593). The AARM fabric
is usually normal (D0923, D0872, D1274, D1309,
D1215) or false normal (D1176) to the AMS.
In the group of the oblate-fabric samples, the HFP
fabric is normal with respect to the AMS and to the
bedding plane (Fig. 9, D1033). The AARM fabric is
normal (D1276, D1253, D1315) or false normal
(D1279, D1118, D1086) to the AMS. In all cases, the
magnetic foliation is subparallel to the bedding plane.
The shapes of AMS, AARM, and HFP ellipsoids are
correlated quantitatively (Fig. 10). The samples with
prolate AMS fabric show triaxial to oblate shapes
compared to their AARM ellipsoids (Fig. 10a). For the
triaxial AMS fabric samples, the shape of the AARM
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
155
Fig. 9. The anisotropy degree vs. ellipsoid shape (P–T) plot of the AMS with equal-area, lower-hemisphere projections of the principal directions of
the AMS (black symbols), high-field paramagnetic anisotropy (gray symbols), and anisotropy of anhysteretic magnetic remanence (open symbols)
tensors in bedding coordinate system. Square, triangle, and circle symbols represent maximum, intermediate, and minimum anisotropy directions,
respectively. Low-field magnetic susceptibility (k), and natural remanent magnetization (M0) value are added for each sample.
ellipsoid varies significantly in the oblate to triaxial
field, where a slight negative correlation can be seen.
The oblate AMS fabric samples possess oblate shapes of
the AARM ellipsoids.
Considering the HFP ellipsoids, the prolate AMS
fabric samples have distinctly prolate shapes of the HFP
ellipsoids (Fig. 10b). The behavior of the triaxial AMS
fabric samples is more complex. Sample D0593
possesses triaxial shape of the HFP ellipsoid falling on
the one-to-one correlation line whereas samples D1058,
and D1362 possess oblate shapes of the HFP ellipsoids.
The oblate AMS fabric sample D1033 has a distinctly
oblate HFP ellipsoid.
5.4. Natural remanent magnetization and AF
demagnetization
In order to isolate the characteristic remanent
magnetization component (ChRM), stepwise AF
demagnetization with measurement of the NRM
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M. Chadima et al. / Tectonophysics 418 (2006) 145–162
Fig. 10. Correlation between the shape of the AMS (TAMS) and AARM (TAARM) ellipsoids (a), and the shape of the AMS (UAMS) and HFP (UHFP)
ellipsoids. The AMS represents the whole rock anisotropy, whereas the AARM or HFP represent the ferromagnetic or paramagnetic anisotropy
components, respectively.
after each demagnetization step was carried out. The
NRM values before any AF treatment, M0, vary
significantly from 6.47 × 10− 5 to 9.42 × 10− 2 A/m
(Fig. 11a). A distinct group of samples with
relatively high M0 (M0 N 5 × 10− 3 A/m) is present,
henceforth denoted as the high-NRM samples. After
the AF treatment, all the samples successfully
demagnetized usually to less than 1% of the original
0 value. As an indirect measure of magnetic
coercivity, the medium destructive field (MDF)–the
AF field at which half the M0 is lost – was
calculated from demagnetization curves for each
sample. It is clearly seen that the majority of
samples demagnetize to half of the M0 in the field
lower than 35 mT (Fig. 11b). The high-NRM samples
demagnetize in the narrow coercivity window between
20 and 30 mT (Fig. 11b). The high MDF (MDF N 35
mT) is bound to the reverse-polarity samples possessing
two antiparallel components of NRM. Due to the normal
to reverse polarity transition, the samples are demagnetized to half of the M0 in relatively high AF fields (Fig.
11b). There is no simple correlation between the M0 and
magnetic susceptibility values (Fig. 11c). The group of
high-susceptibility samples possess relatively low M0
values whereas the high-NRM samples always have
magnetic susceptibility lower than 250 × 10− 6 SI. In the
correlation between the M0 and the quantitative AMS
parameters, the high-NRM samples possess either
oblate or triaxial AMS (Fig. 11d). Despite the even
distribution of the high-NRM samples, it is evident that
all the triaxial-fabric samples are characterized by high
M0 values. On the other hand, the prolate-fabric samples
possess relatively low M0 values. No statistical
dependence was found between M0 and the ratio
between SIRM and magnetic susceptibility (Fig. 11e).
Plotting the S-ratio for a backfield of 300 mT against M0
shows that the high-NRM samples have a slightly higher
S-ratio with an average value of 0.98 × 0.02. However, it
is not greatly different from the remaining samples
which have an average S-ratio of 0.96 × 0.02 (Fig. 11f).
6. Discussion and paleomagnetic implications
As evident from the anisotropy and NRM
measurements, several different groups of sample
behavior can be distinguished. The first group is
represented by the samples with prolate inverse fabric
(Fig. 7d, f) possessing the highest magnetic susceptibility (Fig. 7d), and relatively low M0 values (Fig.
11d). In the studied sections these high-susceptibility
inverse-fabric samples are usually bound to the
concretions and cemented layers. The HFP fabric,
which reflects the paramagnetic component of the
magnetic anisotropy, is parallel to the AMS and
inverse to the bedding plane. On the other hand, the
AARM fabric, which is only due to the ferromagnetic
component of the magnetic anisotropy, is always
normal to the bedding plane (Fig. 9). In the cemented
layers, the iron-bearing carbonate siderite was sometimes macroscopically observed. Siderite is common
in hydrothermal mineralizations but is also encountered in sediments as a result of diagenesis under
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
157
Fig. 11. The natural magnetic remanence (NRM) for the entire set of samples. Frequency distribution of the M0 (a), relationship between M0 and
medium destructive field (MDF) (b), relationship between M0 and magnetic susceptibility (c), and the AMS P–T plot with M0 values of respective
samples plotted as grayscales (d), relationship between M0 and SIRM/k ratio (e), and the relationship between M0 and S-ratio calculated from the
SIRM and IRM acquired in the field of 0.3 T in opposite direction, for selected samples only (f).
158
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
anoxic conditions (Ellwood et al., 1988). Siderite is a
paramagnetic mineral characterized by an inverse
relationship between the crystallographic axes and
the AMS principal axes (Jacobs, 1963). This suggests
that the AMS in the inverse-fabric samples is
controlled by the presence of siderite crystals
distributed with the c-crystallographic axis normal to
the bedding plane. The bedding parallel AARM fabric
reflects the minor presence of a ferromagnetic fraction
distributed within the bedding plane. The presence of
siderite in the prolate-fabric samples was further
verified by XRD analysis (Fig. 3, D1245, Table 2).
During the laboratory heating, the magnetic susceptibility of the siderite-bearing samples decreases up to
300 °C following paramagnetic behavior. Susceptibility resolution into paramagnetic and ferromagnetic
components showed that magnetic susceptibility is
almost entirely carried by paramagnetic component
(Fig. 2, D1245). The rapid susceptibility increase
above the temperature of 300 °C can be attributed to
the laboratory induced mineral changes; siderite is
transformed to maghemite, and later, to magnetite or
hematite (Ellwood et al., 1986; Hirt and Gehring,
1991; Pan et al., 2000) as evidenced by the
susceptibility decrease above approximately 550 °C
(Fig. 2). If the high-susceptibility and the inverseHFP-fabric criteria are applied, it seems probable that
the magnetic anisotropy of the triaxial-fabric sample
D0593 is also siderite controlled (Figs. 9 and 10b). On
the other hand, the HFP fabric of the prolate to triaxial
samples D1058 and D1362 is more or less bedding
controlled (Figs. 9 and 10b). Moreover, these samples
possess relatively low magnetic susceptibility
(k b 152 × 10− 6, Fig. 7d), and relatively high M0 values
(M0 N 4.7 × 10− 2, Fig. 11d).
After AF demagnetization, two stable components
of the NRM can be isolated for the majority of the
siderite-bearing samples (Fig. 12a): a low-field
component (approximately 5–15 mT), and a highfield component (15–80 mT), regarded as ChRM
component. The mean direction of the ChRM
component is D = 46.5°, I = 85.6°, k = 21.6, α95 = 6.3°
(Fig. 12b). This direction is in agreeement with the
magnetic field for the J/K (Savostin et al., 1993). For
all the siderite-bearing samples, only normal magnetic polarity is present (Fig. 12b). Although the
magnetization of the siderite-bearing samples may
have been acquired during the diagenesis of the
sediment, the absence of the reverse polarity
magnetization is suggestive of remagnetization. For
this reason, the siderite-bearing samples were excluded from further magnetostratigraphic study.
The group of the triaxial-fabric samples possesses a
relatively low magnetic susceptibility (Fig. 7d) but the
highest M0 values (Fig. 11d). These high M0 values are
not exclusively found in the triaxial-fabric samples but
can be observed for some oblate-fabric samples as well.
Since the high M0 value is the characteristic feature of
the triaxial-fabric samples, it seems probable that the
high M0 value is not the response to subsequent
magnetization, possibly artificially imparted, but is a
material feature of this particular group of samples. The
high-NRM samples possessing both intermediate and
normal fabric will be considered as one distinct group in
further discussion. The HFP foliation is close to the
bedding plane (Fig. 9, D1058, D1362, D1033) and the
fabric is distinctly oblate (Fig. 10b). The ferromagnetic
fabric, reflected by AARM, is coaxial with the AMS
(Fig. 9) and triaxial in shape (Fig. 10a). As shown by the
linear dependence of the torque moment to the square of
the applied field, the paramagnetic minerals are the
dominant components of the magnetic anisotropy (Fig.
8). It appears that the paramagnetic mineral, which has
normal magnetic fabrics, most probably iron-bearing
chlorite or mica (Fig. 3 and Table 2), is the main carrier
of magnetic anisotropy. The dominance of paramagnetic
mineral(s) on bulk susceptibility can be evidenced by
the magnetic susceptibility versus temperature curve
during laboratory heating up to 240 °C. The curve
follows a paramagnetic behavior and susceptibility
resolution into paramagnetic and ferromagnetic component shows that the paramagnetic contribution to the
room temperature magnetic susceptibility is approximately 80% (Fig. 2, D0653). The two-step increase of
magnetic susceptibility in the temperature interval of
240–400 °C may be attributed (1) to decomposition of
an iron hydroxide or phase transition in pyrrhotite
(Dekkers, 1989; Zapletal, 1993); (2) the decomposition
of berthierine, an iron-rich clay mineral found in
ferruginous rocks (Hirt and Gehring, 1991); or (3) to
the gradual changes in iron sulfide mineralogy (Krs et
al., 1992). Further susceptibility increase can be
attributed to the thermal decomposition of pyrite and
its conversion to magnetite. The abundant presence of
pyrite in the high-NRM samples was evidenced by XRD
(Fig. 3 and Table 2). The pyrite decomposition at about
400–440 °C has been verified by XRD thermal scans
(not presented). The presence of newly created magnetite is verified by a sharp susceptibility decrease at
approximately 550 °C.
In order to explain the ferromagnetic mineralogy in
the high-NRM triaxial-fabric samples, several hypotheses were proposed. An iron sulfide may explain the
high M0 intensities and triaxial-fabrics, but there is no
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
159
Fig. 12. Zijderveld diagrams of the AF demagnetizations for representative test samples (a, c, e), and directions of the characteristic remanence
components for the entire group of samples (b, d, f, gray and open symbols represent normal or reverse polarity, respectively), all in the tilt correction
coordinate system.
evidence that the ferromagnetic mineralogy of these
samples different from the other samples. Susceptibility
did not show any field dependence as would be expected
for pyrrhotite (Fig. 5b), and relatively low SIRM/k ratios
(Fig. 11e), and high S-ratio which is indicative of
presence of the low coercivity mineral, i.e. magnetite
(Fig. 11f) precludes greigite as a carrier of the NRM.
Moreover, the MDF of the high-NRM samples in the
range of 20–30 mT is generally too low with respect to
the values reported for greigite (Dekkers and Schoonen,
1996) or pyrrhotite (Dekkers, 1988).
Another possible explanation of the magnetic
mineralogy of the high-NRM samples may be the
presence of antiferromagnetic goethite which decomposition is suggested by thermal changes of bulk
magnetic susceptibility (Fig. 2). The interference of
the inverse magnetic fabric carried by goethite with the
normal magnetic fabric due to the paramagnetic matrix
was responsible for anomalous magnetic fabric in
several horizons of Pleistocene marine sequence in
central Italy (Lehman et al., 1996). The reported
anomalous magnetic fabric was almost isotropic with a
very low magnitude of anisotropy. This observation is in
correspondence with the triaxial fabric possessing a
relatively low degree of anisotropy (Fig. 9). Despite
these indirect clues the presence of goethite has to be
rejected according to IRM acquisition curves where all
the test samples reached the SIRM at about 0.5 T (Fig.
5a). The Hcr values are also very low compared to the
data published for goethite (Peters and Dekkers, 2003).
Intermediate magnetic fabric for the group of the
triaxial-fabric samples could be possibly explained by
the substrate-controlled growth of siderite crystal along
the bedding plane as described by Hounslow (2001). In
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M. Chadima et al. / Tectonophysics 418 (2006) 145–162
such a situation the c-axes of siderite are oriented
obliquely to the bedding plane resulting in the oblique
magnetic fabric as observed in the siderite-bearing
pelitic rock in SW England (de Wall and Warr, 2004).
Although siderite is abundant in the siderite concretion
and siderite-cemented layer its presence was not found
in the high-NRM samples (Fig. 3 and Table 2). The
magnetic susceptibility of this group of samples is
relatively low as opposed to the siderite-bearing inversefabric samples where the highest susceptibility values
were measured (Fig. 7d). Moreover, the group of
inverse-fabric siderite-bearing samples is rather distinctly delimited (Fig. 7f). It seems probable that the
substrate-controlled growth of siderite crystals is not
the mechanism responsible for the intermediate magnetic fabric in the high-NRM samples.
Despite the above-outlined hypothesis, all the rock
magnetic measurements on the test samples, as well as
SIRM, and the S-ratio values show no significant
difference between the group of normal-fabric and highNRM samples. The normalized IRM acquisition curves,
low Hcr values (Fig. 5a) and high S-ratio (Fig. 11f), are
indicative for Ti-poor magnetite as the main carrier of
NRM in both normal-fabric and high-NRM samples.
After the AF demagnetization, only one stable
component of the NRM can be isolated (5–80 mT) for
the majority of the high-NRM samples (Fig. 12c).
Although both magnetic polarities are present, the
directions are highly scattered (Fig. 12d). Consequently,
the mean direction for samples with normal polarity is
D = 152.0°, I = 78.2°, k = 1.8, α95 = 37.5° and reverse
polarity, D = 320.4°, I = − 61.8°, k = 2.9, α95 = 33.2. Due
to the high directional scatter, the ChRM isolated from
the high-NRM samples cannot be used for intended
magnetostratigraphic research. The magnetic mineralogy of these samples is not different from the other
samples, therefore it is not possible to speculate on why
these samples are not a good recorder of the ancient
field.
The group of the oblate-fabric samples comprises
both relatively low and high-NRM samples (Fig. 11d).
The AMS, HFP, and AARM fabrics are more or less
controlled by the bedding plane (Fig. 9, D1279, D1276,
D1315, D1253, D1118, D1086), whereby the maximum
and intermediate directions of AMS and AARM may be
mutually interchanged. All types of magnetic fabric
(AMS, HFP, AARM) are oblate in shape (Fig. 10a, b). It
appears that the AMS is controlled by the preferred
orientation of the paramagnetic phyllosilicates which
dominate the XRD scans (Fig. 3).
After the AF demagnetization, two stable components of the NRM can be usually isolated (Fig. 12e)
where the high-field component is regarded as the
ChRM. Both magnetic polarities are present for the
ChRM. The mean directions for the normal and reverse
polarities are D = 7.1°, I = 77.9°, k = 21.7, α95 = 2.9° and
D = 276.9°, I = − 68.3°, k = 6.8, α95 = 11.8°, respectively.
The high inclination of normal and reverse mean
directions is in agreement with the magnetic field
direction for the J/K (Fig. 12f).
7. Conclusions
In this study of the Mesozoic black shales on the
northern margin of the Siberian Craton, we have shown
that the low-field magnetic susceptibility is predominately carried by the paramagnetic minerals, i.e. iron
carbonate siderite, iron-bearing chlorites, and micas.
Despite this fact, the influence of ferromagnetic fraction
on the magnetic properties of some samples cannot be
neglected. According to the magnetic anisotropy
behavior and NRM, three main groups of samples can
be distinguished: i) siderite-bearing samples that are
characterized by an inverse prolate magnetic fabrics,
usually accompanied by relatively high magnetic
susceptibility values (k N 300 × 10− 6 SI), ii) a triaxialto oblate-fabric samples possessing high M0 intensities
that are one to two orders of magnitude higher than
average M0, and iii) oblate-fabric samples with a
relatively low M0 values. There is no apparent
correlation between the high susceptibility and high
M0 values; on the contrary, the high M0 excludes the
high susceptibility, and vice versa.
The low-field magnetic anisotropy in the sideritebearing rocks (i) is controlled by the preferred
orientation of siderite known for its inverse magnetic
anisotropy. The ferromagnetic anisotropy, determined
by the AARM, is oblate in shape and controlled by
bedding compaction. Two stable components of the
NRM can be isolated for the majority of the sideritebearing samples. The high-field component, regarded as
the ChRM, is characterized by absence of reversepolarity samples, suggesting that this component is a
secondary magnetization, most likely of chemical
origin. For this reason, the siderite-bearing samples
will be excluded from the magnetostratigraphic
investigation.
The magnetic mineralogy of the high-NRM, triaxialto oblate-fabric samples (ii) is more complex. The AMS
is mainly controlled by the paramagnetic, and to the
lesser extent by ferromagnetic fraction. Ferromagnetic
minerals are probably of detrital origin. Only one stable
component of the NRM can be isolated for the majority
of the high-NRM samples. Although both magnetic
M. Chadima et al. / Tectonophysics 418 (2006) 145–162
polarities of the ChRM are present, the directions are
highly scattered. The anomalous AMS results and the
scattered ChRM directions also suggest that the
magnetic mineralogy has undergone some change in
the past. These samples will also be excluded from the
magnetostratigraphic study.
The magnetic anisotropy in the group of the
oblate-fabric samples (iii) is predominantly controlled
by the preferred orientation of iron-bearing chlorites
or micas, and to a minor extent, by the ferromagnetic
fraction. Two stable components of magnetization
can be usually isolated. Both magnetic polarities are
present for the high coercivity ChRM. The oblate,
bedding-controlled fabric, and the low M0 suggests
that these samples may be good recorders of the
ancient field and therefore useful for magnetostratigraphic study across J/K boundary. This group
represents approximately 85% of the entire sample
set, and covers evenly the entire length of the studied
section.
Acknowledgements
The authors gratefully acknowledge the colleagues
with whom they made a field party in the tough arctic
conditions in August 2003: V. Zakharov, M. Rogov, M.
Košt'ák, and M. Mazuch. Laboratory measurements and
some data processing were carried out by N. Nowaczyk,
P. Vorm, P. Bezdička, V. Goliáš, Z. Korbelová, A.
Langrová, and J. Petráček. The comments and suggestions from L. Sagnotti and anonymous reviewer and
discussion with E. Petrovský considerably improved the
manuscript. The attendance at the AGU 2004 Fall
Meeting was supported by Hlávka Foundation, Prague.
The research was supported by the Czech Science
Foundation Research Grant #205/02/1576.
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