RUSSIAN JOURNAL OF EARTH SCIENCES VOL. 8, ES3001, doi:10.2205/2006ES000204, 2006

Results of Petromagnetic Measurements

Specific magnetic susceptibility (χ), specific saturation magnetization (Ms), and specific saturation remanent magnetization (Mrs) (Table 1)

2006ES000204-fig03
Figure 3
[20]  The values of these characteristics vary within wide limits, generally reflecting the main lithologic properties of the rocks such as the contributions of diamagnetic material (calcite and quartz), paramagnetic material (Fe-bearing clays and Fe hydroxides), and magnetic minerals of terrigenous origin; accordingly, the magnetization is lowest in Maestrichtian marls and Danian interbeds K, S, and T enriched in diamagnetic calcite and quartz, whereas the sandy-clayey sediments of the upper part of section (layers U, V, and W) are most magnetic. Both groups of sediments are largely affected by paramagnetic material whose magnetization is about 10-20 times higher than the saturation magnetization (  Ms ) of magnetic minerals (Table 1), and the amount of paramagnetic material in the sandy-clayey deposits is about three times larger compared to marls and limestones (Table 1). The positive correlation between Ms and Mrs (Figure 3) implies a decisive role of both concentrations of magnetic minerals. The correlation of Ms and Mrs
2006ES000204-fig04
Figure 4
with the magnetic susceptibility is less distinct (Figure 4). Apparently, the susceptibility is appreciably affected by the contributions of paramagnetic, diamagnetic (divergences in the weakly magnetic region), and superparamagnetic (divergences in the strongly magnetic region) materials; these effects are largely eliminated from Ms and are absent in Mrs. The "divergences" in Ms and Mrs in sample W can be accounted for by the presence of numerous fine magnetic grains making a small contribution to Ms.

2006ES000204-fig05
Figure 5
2006ES000204-fig06
Figure 6
2006ES000204-fig07
Figure 7
TMA data
[21]  (Table 2). The analysis of the curves M(T) (Figure 5) and their derivatives (Figure 6) has identified seven magnetic phases.

[22]  (1)  TC = 90-150oC, the phase accounts for 10-20% Ms. It is present in all samples studied (Table 2) and is destroyed upon heating (Figure 5). Most likely, it consists of ferromagnetic iron hydroxides of the goethite type. Assuming that this is goethite with Ms = 0.02 A m2 kg-1, we obtain that its concentration varies in the section from 0.2-0.5% in marls of the Maestrichtian and in the lens K and interbeds S and T of the Danian to 2-3% in the sandy-clayey sediments (Figure 7a). The bulk concentration of iron (Fe2O3 ) in the deposits varies, respectively, from 2% to 6-8% [Grachev et al., 2005]. Therefore, the amount of paramagnetic iron varies from ~1.5 to 3-4%. This is the iron of paramagnetic iron hydroxides and/or iron-bearing clayey minerals.

2006ES000204-fig08
Figure 8
[23]  (2)  TC = 180-300oC, the phase is present in all samples except the boundary layer J (after heating, this phase also arises in J samples). It accounts for 5-40% Ms (Table 2). After heating to 800oC, its fraction in many samples increases by 30-90%, and the Curie point generally decreases (Figure 5 and Table 2). Successive heatings of samples (e.g. sample K, Figure 8) reveal that this rise takes place only after heating to 800oC. An increase in Ms associated with a drop in TC implies that this is hemoilmenite partially homogenized during heating; as a result, the curve Ms(T) is typically concave. Check heatings of some samples to 1000oC showed that the concavity of the curve Ms(T) disappears, and the value Ms is noticeably larger compared to the results of heating to 800oC (Figure 2). This corresponds to the state diagram of hemoilmenite of an intermediate composition for which the region of the homogeneous state lies above 900oC [Nagata, 1961]. Results of the second heating were used to determine the hemoilmenite concentration (from the TC versus hemoilmenite composition diagram [Nagata, 1961]). This concentration varies from less than 0.0001% to 0.02% (Figure 7b). The magnetic fractions of all samples studied with the microprobe contain large numbers of ilmenite grains (clasts); the latter are often well-preserved relatively large plates more than 50  m
2006ES000204-fig09
Figure 9
m in size (e.g. see Figures 9a, b). Their concentrations in the sediments amount to a few tenths percent. The composition is close to pure ilmenite (Table 3). They often contain intergrowths and lamellae of rutile. No hemoilmenite grains were observed whose composition corresponds to a Curie point of 200-300oC. The majority of the hemoilmenite grains are very fine (smaller than the probe size), as is evident, for example, from their high coercivity (see below), and the concentration of hemoilmenite with TC = 180-300oC is one to two orders lower than the concentration of ilmenite. Very thin lamellae of hemoilmenite (tenths and hundredths of micron) are poorly observable in the ilmenite grains, and their composition could not be measured by the microprobe 2-3  m m in size. Moreover, ilmenite is well drawn away by a magnet, apparently, due to hemoilmenite inclusions.

2006ES000204-fig10
Figure 10
[24]  It is possible that Mg-Al-ferrospinels with similar Curie points (200-300oC) could form during laboratory heatings. In this case, rocks must contain silicates containing iron, magnesium, and aluminum and decomposing at high temperatures (e.g. see [Bagin et al., 1976, 1977; Gapeev and Tsel'movich., 1988]). The study rocks (particularly, sandy-clayey sediments) contain a sufficient amount of components necessary for such a process [Grachev et al., 2005]. However, no correlation of the amount of this magnetic phase with Fe, Mg, and Al concentrations is observed. For example, in samples from the layer J, containing the highest concentrations of the above elements (7-8% Fe2O3, 17-19% Al2O3, and 2.6-3% MgO), the magnetic phase with TC = 200-300oC virtually does not form during successive heatings: the curves M(T) nearly coincide up to 850oC (Figure 8), whereas the concentration of this magnetic phase obviously rises after heating above 800oC in a sample from the lens K (Figure 8), in which the concentrations of the above elements are much lower (3.5-4.7%Fe2O3, 5.7-8.9%Al2O3, and 0.9-1.3%MgO). Another argument is provided by the thermomagnetic study of the magnetic fraction extracted by a permanent magnet from samples of the layers L and W and their "nonmagnetic" residues. As seen from Figure 10a, a Curie point of about 250oC is fixed precisely in the magnetic fraction; the relative amount of the latter is small (less than 20%), and the curves M(T) of the second and third heatings lie appreciably lower than the first heating curve, which may be caused by destruction of nearly half of magnetite possibly due to its heating-related oxidation. One might expect that Mg-Al-ferrospinels would form most intensely from the nonmagnetic fraction, but this is not observed: its heatings caused no alterations (Figure 10b).

[25]  (3)  Tb = 340-370oC, the phase is observed in all samples of the layer J and in samples of the layers R and T. As seen from the data of successive heatings, this phase is generally destroyed after heating to 300oC (Table 2, Figure 8); i.e. in the vast majority of cases, this is not a Curie point but a result of destruction of a magnetic mineral. Most likely, this is the usual process of the transformation of maghemite into hematite.

2006ES000204-fig11
Figure 11
[26]  The thermomagnetic and microprobe examination of sample J6 and its magnetic fraction revealed metallic nickel with a Curie point of about 360oC in two pieces less than 3 mm in size from the upper (sample J6-6) and the middle (J6-4) parts of the layer J (for more details, see [Grachev et al., 2005]) and in samples J2 and J3. The curves M(T) from the remaining samples, including samples J6-1, 2, 3, and 5 and even a small piece taken near sample J6-6, yield no evidence for metallic nickel (Table 2), but the latter is detected from the curves Mr(T) of several samples (Figure 11), although there are samples (e.g. J4-1) that do not contain nickel from Mr(T) as well (Figure 11). These results imply that, first, nickel exists as very fine grains whose average concentration in the layer J is apparently less than 0.001 (~0.02, ~0.01, and ~0.1% in pieces from samples J3-2, J6-4, and J6-6, respectively); therefore, they have no signature in the value of Ms but contribute to Mrs. Second, the detection of metallic nickel only in individual minute samples is evidence for its local and very irregular distribution in the layer J. Apart from the layer J, an intergrowth of pure nickel and copper was discovered in sample L6
2006ES000204-fig12
Figure 12
(Figure 12). Thermomagnetic analysis have not discovered nickel in the layer L, supporting its very irregular distribution. The presence of individual Ni grains in the layer L is possibly due to the erosion of the upper part of the layer J and the redeposition of Ni particles that settled mainly during the deposition of the upper part of the layer J.

[27]  (4)  TC = 550-610oC, the phase is present in all of the studied samples of the section and accounts for 20% to 60% of Ms (Table 2 and Figure 5). After heating, this phase is generally preserved, but its amount usually decreases and TC in several samples shifts to the left. Only in two cases, in samples K and T, the amount of magnetite increases after heating (Table 2 and Figure 5). Often this is titanomagnetite successively oxidized to magnetite; in turn, the latter is often single-phase oxidized ( TC>580o C). After a fast laboratory heating to 800oC, titanomagnetite grains are partially homogenized. This feature implies the presence of titanomagnetite in samples from the Maestrichtian layers B, C, E, G, and H and from the Danian layers J, R, V, and W. In the other layers, titanomagnetite is absent, but magnetite is present; this is valid for the layers L, M, N, and U, distinguished by the highest concentrations of magnetic minerals. The absence of titanomagnetite is confirmed by the microprobe data: solely magnetite that does not contain titanium is discovered in the layers K, L, M, O, and P (Table 3). The magnetic fraction from the layer W includes very fine grains dominated by ilmenite (Table 3), and titanomagnetite is fixed only from TMA data. The presence of titanomagnetite in the layer J is confirmed by microprobe data: the composition of grains is close to titanomagnetites typical of basalts (TiO 2sim 20-25%) [Grachev et al., 2005]. Approximate estimates of the magnetite and titanomagnetite concentrations in the samples vary from < 0.0001% to 0.001% (Figure 7c). Moreover, the presence or the absence of titanomagnetite and its concentration correlates in no way with the concentrations of magnetite, i.e. they have different sources. Magnetite clasts of the inspected magnetic fractions very often contain well-preserved single crystals (octahedral, Figures 9c, d), which is evidence for a near source area or in situ crystallization of magnetite. Such crystals of pure magnetite are evidently of nonmagmatic origin.

[28]  (5)  TC = 640-660oC, the phase is present only in samples from the layer J (Table 2) and accounts for 10-15% Ms. After heating to 800oC, this phase is destroyed, implying that this is not hematite. Taking into account the presence of nickel in samples from the layer J, we may suppose that this is a Fe-Ni alloy, and a simple calculation of TC and Ms for iron and nickel shows that this can be Fe3Ni, which is confirmed by the microprobe data [Grachev et al., 2005].

[29]  (6)  TC = 660-670oC, the phase is only present in a sample from the lens K and is preserved after heating to 800oC. It is evidently hematite. After heating to 800oC, hematite forms in half of the samples studied and has TC = 660-680oC (Table 2).

2006ES000204-fig13
Figure 13
[30]  (7)  TC = 740-770oC, the phase is present in 19 samples and its contribution to Ms amounts to 10-30% (Table 2, Figures 5 and 6). After heating to 800oC, this phase is partially or completely destroyed (Figure 5). Evidently, this is fine grains of metallic iron with minor admixtures that oxidizes during heating to 800oC. Individual balls of pure iron were discovered by the microprobe in samples J2 and M4 (Figure 13). Its concentration is small, less than 0.0006%. In the layer J, the thermomagnetic analysis did not discover metallic iron, but a magnetic species with TC = 640-660oC is present in the layer (see above); probably, it is a Fe-Ni alloy with a concentration of no more than 0.0002%. The along-section distribution of metallic iron (including its alloy with nickel) is rather uniform (Figure 7d).

2006ES000204-fig14
Figure 14
Coercivity of magnetic minerals and coercivity spectra.
[31]  As seen from coercivity spectra (CS) presented in Figure 14, all samples have similar ensembles of magnetic grains. The CS of the Maestrichtian marls are least different and very close to the CS of the layers K, S, and T in the Danian deposits. The spectra exhibit a smooth increase to a maximum at 100-140 mT and a subsequent drop to a minimum at ~400 mT followed by a rise until a field of 500 mT. The CS extrema in the sandy-clayey deposits virtually disappear beginning from the layer L: the CS smoothly increase until a limiting field of 800 mT used in the measurements (Figure 14). In the upward direction along the section, the CS are gradually transformed into marl-similar CS: in the interval from samples M to S and T, a plateau first appears and, in the overlying layers, it is transformed into a distinct maximum at 130-160 mT and a minimum at ~400 mT. The CS of the uppermost horizons of the section are similar to those of the Maestrichtian marls (Figure 14).

[32]  Against this background, the CS of the layer J is markedly distinguished in its low coercivity region by a maximum or a plateau at 25-40 mT. In the remaining part, the CS of the layer J is similar to CS from samples of sandy-clayey deposits, particularly in the layers N and O (Figure 14).

2006ES000204-fig15
Figure 15
[33]  The CS being stretched, its integral characteristics such as HCr and Mrs/Ms are smoothed and its dependences on the compositions of rocks and minerals are averaged (Table 1), but a decrease of coercivity is seen in the HCr of the layer J. Judging from the values of HCr and Mrs/Ms (Table 1 and Figure 15), single-domain (SD) and pseudosingle-domain (PSD) magnetic grains prevail in the rocks,
2006ES000204-fig16
Figure 16
but the vast majority of points in Day plot (Figure 16) lie in the multidomain region, which is due to the presence of a large number of superparamagnetic grains. In high fields, their Ms effect can be eliminated together with the paramagnetic effect (if the superparamagnetic grains are very fine). However, in a lower field of the order of HC, the susceptibility of these grains is high and, for this reason, remagnetization takes place in a field much smaller than the real HC value. This leads to overestimation of the ratio HCr/HC. The superparamagnetic magnetization curve is not linear, as in the paramagnetic case (at room temperature), but hyperbolic, typical of ferri- and ferromagnetic species. Eliminating the paramagnetic magnetization through a linear approximation, we do not remove the superparamagnetic contribution in MsMrs is not influenced by the superparamagnetism). This decreases the ratio Mrs/Ms. As a result, points in the Day plot are displaced to the right and downward. Overall, notwithstanding the distortion of concrete values, the HCr/HC-Mrs/Ms diagram displays a general tendency (Figure 16). As seen from the Day plot, the finest SD magnetic grains are present in the layers K and T (Table 1 and Figure 16). The low coercivity of magnetic grains in the layer J has no effect on Mrs/Ms and HCr/HC, thereby emphasizing that these ratios are unrelated to the size of magnetic grains.

2006ES000204-fig17
Figure 17
[34]  Figure 17 presents the magnetization curves (up to fields of 100 mT) of superparamagnetic particles in samples studied. For clearness, the curves are normalized to the maximum superparamagnetic magnetization. They can be used to establish which superparamagnetic grains, coarse or fine, prevail in the spectrum. Very rapid saturation is evidence for the presence of coarse particles, and prolonged saturation indicates fine grains. On the other hand, rapid saturation implies the presence of grains with large values of Ms. Our examples show that samples from the layer J are saturated much more rapidly than the remaining samples, which means that larger and probably more magnetic grains are present in the samples J. We consider them as nickel grains. The slowest saturation is observed in samples of sandy-clayey rocks, i.e. they contain the finest and probably the least superparamagnetic grains (e.g. goethite and hemoilmenite whose concentrations in the sandy-clayey deposits is appreciable higher compared to marls (Figures 7 and 17)).

2006ES000204-fig18
Figure 18
[35]  Now, we consider the correlation of such CS characteristics as the position and height of main extrema with the concentration of magnetic minerals (Figure 18). Such a correlation is absent in the case of metallic iron, which can be attributed to its small concentration. We relate a weak correlation for "goethite" to the presence of a complex association of iron hydroxides that includes, high coercivity grains of needle goethite, its predominantly low coercivity earthy varieties, and paramagnetic hydroxides of iron. A positive correlation with magnetite+titanomagnetite and hemoilmenite is observed, particularly, in sandy-clayey rocks (Figure 18). This suggests that the CS of the studied rocks is mainly controlled by grains of magnetite, titanomagnetite, and hemoilmenite. We emphasize that the ilmenite grains that are present in noticeable amounts in the sandy-clayey sediments are paramagnetic at room temperature, i.e. they do not contribute to CS. In the layer J, most likely, metallic nickel and its alloy with iron, make the main contribution to the low coercivity part of the spectrum.

Anisotropy.
[36]  We measured the anisotropy of the magnetic susceptibility A<undef>c and the saturation remanent magnetization Ars. The first involves all minerals, magnetic, superparamagnetic, paramagnetic, and diamagnetic, while the second is related solely to magnetic minerals. On the whole, both types of anisotropy behave, with rare exceptions, similarly (Table 1). The Ac values of the main group lie within the limits 1-1.1 and only four samples yielded Ac>1.1; whereas the main part of Ars varies from 1.12 to 1.36, and only four samples yielded Ac le 1.11. Apparently, this is due to the fact that the paramagnetic and diamagnetic parts of the sediments are, on the whole, isotropic, although calcite and clayey minerals are anisotropic ( Ac is 1.13 in calcite and 1.2-1.35 in clays, whereas quartz is isotropic [Rochette et al., 1992]) and the distribution of their symmetry axes in the studied sediments is close to chaotic. Therefore, the magnetic susceptibility anisotropy is determined in our case by magnetic minerals. The Ars values show that, with rare exceptions (samples K and T), the studied sediments are anisotropic and the anisotropy depends weakly on the composition of the rocks. In the interval from A to R, the anisotropy of Ars varies within close limits but is appreciably enhanced in the upper horizons U-W. Within each layer, Ars varies within narrow limits except for the boundary clay J, where the anisotropic scatter is widest, from 1.02 to 1.32.

[37]  In the vast majority, the sediments of the section possess a foliation fabric ( E>1 ) and only some of its beds are characterized by either Esim 1 or a very weak lineation ( E<1, samples from the Maestrichtian layers B, C, G, and H) (Table 1). All of the aforesaid can be accounted for by the presence of elongated grains of magnetic minerals, compaction of the sediments, a certain influence of currents, and so on. The presence of anisotropy and a magnetic fabric is evidence for a terrigenous origin of magnetic minerals that are the main carriers of magnetization in the sediments. Authigenic magnetic minerals are likely present in isotropic samples (K, T, and others). Apart from the normal magnetic fabric (the minimum susceptibility is perpendicular to the bed plane), intervals of the inverse fabric (the maximum susceptibility is perpendicular to the bed plane) are also identified. The latter are the beds I, L-Q, V, and W, although they contain normal fabric samples as well (Table 1). Such an inverse magnetic fabric is characteristic of siderite and other Fe carbonates, with their easy magnetization axis being perpendicular to the c symmetry axis [Rochette et al., 1992]. However, appreciable amounts of siderite and the like are not observed in the sediments studied; moreover, as noted above, the paramagnetic and diamagnetic parts of Ac are, rather, isotropic. Therefore, the inverse fabric is more likely related to magnetic minerals. The normal and inverse magnetic fabrics determined from the susceptibility and remanent magnetization coincide, which additionally confirms the conclusion on the noticeable contribution of magnetic minerals to the magnetic susceptibility (Table 1). The inverse fabric determined from the magnetic susceptibility is known for needle goethite and elongated (uniaxial) SD grains of magnetite. In both cases, the susceptibility is minimal along the longer axis of the grain, i.e. the easy magnetization axis is perpendicular to the elongation direction of the grain [Rochette et al., 1992]; the same is true of the remanence of needle goethite [Bagin et al., 1988]. Inverse fabrics of rocks determined from remanent magnetization have repeatedly been observed and are often related to a tectonic factor [Rochette et al., 1992]. In our case, the undeformed state of sediments of the sequence excludes a tectonic factor as the cause of the inverse magnetic fabric. The amount of magnetic anisotropy and the characteristics of the magnetic fabric do not correlate with the composition and concentration of magnetic minerals, but the following general tendency can be noted: the magnetic fabric is invariably normal in marls in which the concentration of goethite is much smaller compared to the sandy-clayey sediments (Tables 1 and 2); therefore, it is likely that the inverse fabric is primarily related to the presence of needle goethite in sediments.

2006ES000204-fig19
Figure 19
Behavior of magnetic properties along the section.
[38]  Two levels of c, Mrs, and Ms are clearly fixed in the section: (1) weakly magnetic Maestrichtian marls underlying the layer J, the lens K, and the interbeds S and T in Danian sediments; and (2) more magnetic sandy-clayey sediments of the layers J and L-W (Table 1, Figure 19). These two levels are generally recognizable in the along-section distributions of magnetite, hemoilmenite, and goethite (Figures 7a, 7b, and 7c) but are absent in the distributions of titanomagnetite (Figure 7c) and metallic iron (Figure 7d). The noticeable distinction in the along-section behavior of c, on the one hand, and Mrs and Ms, on the other hand, is evidently due to an appreciable contribution of paramagnetic material and fine superparamagnetic grains to the susceptibility. Within these levels, we can see fluctuations in the magnetization closely correlating with variations in the concentration of hemoilmenite, magnetite, and goethite (relative maximums at - 14-12, 4, and 22-26 cm), implying a certain cyclicity in the accumulation of magnetic minerals; the same pattern is observed in minimums of c, Mrs, and Ms ( - 16, 2, and 18-20 cm) correlating with the anomalous layers K, S, and T. The cyclicity "wave" is most pronounced in the behavior of Mrs (Figure 19c). The layer J does not differ in the overall concentration of magnetic minerals from other levels (Figure 19), but it is distinguished by the lowest coercivity (Figure 15). We relate the latter circumstance to the presence of nickel and an iron-nickel alloy in the layer J. The concentration of magnetic minerals is lowest in the layers K, S, and T, containing predominantly SD magnetic grains; i.e. the cyclicity in the accumulation of magnetic minerals is also expressed in the mean size of their grains (Figure 15). Moreover, samples from the layers K, S, and T differ from the remaining samples by isotropy and a substantial increase in the amount of secondary magnetite due to laboratory heating (Table 2). Therefore, they contain some authigenic (isotropic) magnetic and paramagnetic minerals (e.g. pyrite) that are oxidized during heating and produce magnetite. Specific features of the layers K, S, and T emphasize the cyclicity of the sedimentation process.

[39]  The along-section distributions of goethite, hemoilmenite, and magnetite are generally similar, which implies concurrent accumulation of these minerals under the lithologic control. The chaotic distribution of metallic iron is unrelated to both lithologic properties of the sequence and the K/T boundary (Figure 7d). Titanomagnetite is present at nearly all levels of the Maestrichtian deposits and in the layer J, whereas it occurs only in the upper part of the section in the Danian sandy-clayey sediments (samples R, V, and W). Unlike magnetite, goethite, and hemoilmenite, the concentration of titanomagnetite, wherever it is present, varies insignificantly at all levels. We may state that the presence of titanomagnetite is independent of the lithology of the sequence; rather, taking into account its composition typical of basalts, it characterizes volcanic eruptive activity and the dispersal of titanomagnetite by air. The magnetite concentration is controlled by lithology, although with a certain lag: it is very low (occasionally vanishing) in the Maestrichtian marls up to the layer K (including the layer J) and, only beginning from the layer L, the magnetite concentration increases by about an order (Figure 7c). The hemoilmenite grains exhibit a similar pattern: the hemoilmenite concentration increases substantially (by more than five times) above the layer J (Figure 7b). Lithologic control is most pronounced in the goethite accumulation: an abrupt increase in its concentration is observed precisely in the layer J (Figure 7a).

[40]  Lithologic control is also traceable in values of Ms near 800oC, where the contribution of magnetic minerals vanishes and, accordingly, one may gain constraints on the relative paramagnetic ( Mp ) and diamagnetic ( Md ) fractions in the magnetization of the sediments (Table 1). Overall, the values of Ms at 800oC are positive, i.e. paramagnetic, in the sandy-clayey part of the section and negative, i.e. diamagnetic, in the limestones. More specifically, we may speak of relative paramagnetic and diamagnetic fractions, because noticeable amounts of diamagnetic carbonates and quartz can be present in the sandy-clayey beds, and the same is true of paramagnetic clayey minerals and iron hydroxides in the marls. Thus, the Maestrichtian marls contain paramagnetic material, as is seen from the small values Mp = (1-2) times 10-5 A m2 kg-1 (samples B and G, Table 1) and the small value Md = - 2 times 10-5 A m2 kg-1 in the other Maestrichtian marls. The magnetic susceptibility of paramagnetic minerals is 30 to 300 times higher than the susceptibility of diamagnetic materials [Rochette et al., 1992]. Accordingly, given such small values of Md and Mp, the ratio between the paramagnetic and diamagnetic materials in the rocks under consideration should be at least 1/30; i.e. if, for example, about 2% Fe2O3 were present in marls, the coinciding values of Md and Mp would require more than 60% of diamagnetic calcite and/or quartz. Pure diamagnetic chalk from the Koshak section has Md = - (26-35) times 10-5 A m2 kg-1 [Pechersky et al., 2006]. This value gives an idea of the significance of the paramagnetic admixture in the Gams deposits. The similar values of Md and Mp in the Maestrichtian marls indicate the homogeneity of these rocks. These characteristics vary from +6 to - 12 times 10-5 A m2 kg-1 in the lens K and the Danian interbeds S and T, whereas we have Mp = (15-36) times 10-5 A m2 kg-1 in the sandy-clayey sediments. The layer J differs little in this lithologic indicator: Mp = (26-36) times 10-5 A m2 kg-1.

2006ES000204-fig20
Figure 20
[41]  Now, we compare the behavior of the susceptibility (Figure 19a), saturation magnetization (Figure 19b), and concentration of magnetic minerals (Figure 7) with the behavior of Md and Mp and the bulk concentration of iron, the main magnetization carrier in the rocks (Figure 20). As seen from the comparison between these figures, the along-section distributions of susceptibility and paramagnetic magnetizations at room temperature and at 800oC agree best with each other and with the Fe2O3 concentration. It is clearly seen that the regimes of Fe accumulation in the Maestrichtian and Danian parts of the sequence, fairly homogeneous in each of the parts, are different and rhythmic variations in the sedimentation conditions lead to a decrease in iron in the layers K, S, and T. This cyclicity is recognizable in the accumulation of both magnetic minerals and paramagnetic iron. This correlation is weaker in the behavior of Ms and the concentration of magnetic minerals. This can be due to the fact that more than half of iron in the deposits is present in the paramagnetic form. Thus, the total concentration of the iron oxide in goethite+magnetite+titanomagnetite+hemoilmenite does not exceed 3%, whereas the concentration of Fe2O3 in the sandy-clayey sediments varies from 6% to 8%; accordingly, a half of iron is concentrated in paramagnetic hydroxides of iron and clayey minerals, which differed in the accumulation regime from magnetite and hemoilmenite (apparently, of volcanic-terrigenous origin) and, even to a greater extent, from titanomagnetite (apparently, of volcanic origin).


RJES

Citation: Pechersky, D. M., A. F. Grachev, D. K. Nourgaliev, V. A. Tsel'movich, and Z. V. Sharonova (2006), Magnetolithologic and Magnetomineralogical Characteristics of Deposits at the Mesozoic/Cenozoic Boundary: Gams Section (Austria), Russ. J. Earth Sci., 8, ES3001, doi:10.2205/2006ES000204.

Copyright 2006 by the Russian Journal of Earth Sciences

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