Identification of Ferromagnetic Minerals

IDENTIFICATION OF FERROMAGNETIC MINERALS

MAGNETIC MINERAL ANALYSIS OF PANNONIAN, SARMATIAN AND BADENI AN (10.5-13.7 MA) SEDIMENTS FROM THE SPANNBERG 21 WEL L (EBENTHAL, AUSTRIA)

BSC - THESIS

G.J.H.M. KOOLEN D E L F T , FEBRUARI ‟10

ABSTRACT

The magnetic mineral content of marine and fluviatile sediments in the Vienna Basin, Austria has been measured and described. The magnetic signal measured in a well records the earth‟s magnetic field during deposition of the sediments. Two main ferromagnetic minerals, magnetite and greigite, are found as the dominant natural remanent magnetism carriers. In this study magnetite bearing samples were distinguished from greigite bearing samples. In more than 75% of the samples it was legitimate to decide upon IRM and Curie measurements whether the sample consisted mainly of magnetite or greigite. For the uncertain part of the samples additional experiments were carried out to further determine the mineral content. One of the data processing steps, the construction of IRM components analysis graphs, was rather subjective. Finally it was found that at certain depths the natural remanent magnetism is small because of a small magnetic mineral content. Later formed authigenic greigite is able to lower the overall remanent magnetization due to having an opposite polarity than the polarity being present in the earlier deposited magnetic minerals.

ACKNOWLEDGEMENTS

I would like to thank my academic supervisor of this work ir. W.E. Paulissen, Department of Geotechnology, Delft University of Technology, Delft for her kind supervision, for providing sufficient information required for this thesis and especially for her faith in me and my BSc- project.

Special thanks also go to Dr. M.J. Dekkers, Paleomagnetic Laboratory, Fort Hoofddijk, University Utrecht, Utrecht for his scientific support, for the experiments he carried out on behalf of this BSc-thesis and for his time to explain the theoretical backgrounds of all the magnetic matter discussed. I would furthermore like to thank the Fort Hoofddijk section as a whole for the ability of carrying out all the necessary experiments.

Furthermore I would like to thank Prof. Dr. S.M. Luthi , Head of Department of Geotechnology, Delft University of Technology, Delft for his approval of this thesis.

CONTENT

Abstract ...... 2 Acknowledgements ...... 3 Content...... 4 Introduction ...... 5 2. Geology of the Vienna Basin ...... 6 2.1 Tectonic setting ...... 6 2.2 Sedimentological processes ...... 7 3. Paleomagnetic fundamentals ...... 9 3.1 Basic defenitions in paleomagnetism ...... 9 3.2 The geomagnetic field ...... 9 3.3 Magnetism in rocks ...... 10 3.3.1 Diamagnetism ...... 10 3.3.2 Paramagnetism ...... 10 3.3.3 Ferromagnetism ...... 10 3.4 Magnetic mineralogy ...... 11 3.5 Natural remanent magnetism ...... 12 4. Methods ...... 14 4.1 Samples ...... 14 4.2 Susceptibility measurement ...... 14 4.3 Alternating field demagnetization...... 15 4.4 Isothermal remanent magnetization (IRM) ...... 16 4.5 Anhysteric remanent magnetization (ARM) ...... 16 4.6 Curie temperature loops ...... 17 5. Results ...... 18 5.1 Susceptibility measurement ...... 18 5.2 Curie Temperature Measurement ...... 18 5.3 IRM Component Analysis ...... 23 5.4 Correlation of Magnetization ...... 27 6. Discussion & Conclusion ...... 32 References ...... 34 Appendix ...... 36

INTRODUCTION

The Geological High-resolution Magnetic Tool (GHMT) is a fairly new tool used to measure the remanent magnetization in sediments. The first measurements were carried out in the early 1990‟s. The tool measures the total magnetic field and susceptibility and is dependent on the total magnetic mineral content of sediments. From these two measurements the remanent magnetization can be determined. When the Spannberg 21 well was logged in the Vienna Basin a large variance in the intensity of the signal was found. At certain depths a large GHMT signal was measured whereas at other depths there was a very low signal. The purpose of this BSc-thesis study is to identify the magnetic mineral content of the sediments present in the Spannberg 21 well and to possibly find an explanation for the large variations in the GHMT signal. In total 32 samples at different depths were measured using IRM, ARM, KLY-2 susceptibility and Curie-temperature measurement techniques. AF demagnetizing was first used to demagnetize all the samples. The Curie-temperature measurements were only carried out on 7 samples since they were used to verify the IRM results. All measurements were carried out at the Fort Hoofddijk, Utrecht University, Utrecht. Finally, after sufficient processing, a good identification of the main magnetic minerals (magnetite or greigite) was found for each sample and an explanation is given for the large GHMT signal variations.

G.J.H.M. Koolen

2. GEOLOGY OF THE VIENNA BASIN

2.1 TECTONIC SETTING The Spannberg 21 well was drilled in the Central Paratethyan Vienna Basin. This basin spreads from the Czech and Slovak Republic in the North to northeast Austria in the South. It has a rhomboidal shape and is about 200 km long and up to 55 km wide. The formation of the basin can be subdivided into 4 different geological stages starting from the Lower Miocene and continuing into the present.

Figure 2-1: Four cycles in the formation of the Vienna Basin (modified after Kovac, 2000)

Piggyback basin (Lower Miocene) The formation of the Vienna Basin started in the Early Miocene as an E-W trending piggyback basin on top of the Alpine thrust belt (Figure 2-1 A)(e.g. Decker 1996). The paleostress field was characterized by a NW-SE oriented main compression (Kovac et al. 1989). In this time, sediment deposition was concentrated in piggy-back basins on the folding wedge of the Outer Carpathians and in wrench-fault furrows on the colliding margin of the Central Western Carpathians (Kovac & Barath 1995). The now slowly subsiding sedimentary area was E-W oriented while NW-SE oriented thrusts dominated in the wedge (Kovac 2004).

Pull-apart basin (Middle to Upper Miocene) In the Late Carpathian, thrusting developed into lateral extrusion. This caused a geometric change from a piggyback basin into a rhombic-shaped pull-apart basin (Royden 1985). In the Southern part of the Vienna Basin sedimentation started discordantly with the deposition of the Aderklaa Conglomerate in a braided river system during the Early Badenian sea-level lowstand (Weissenbäck 1996). In the central Vienna Basin the Badenian (16.3 – 12.7 Ma) sediments were divided into proximal deltaic clastics and a distal basinal facies, which is characterized by sandy marls and clay. Carbonates were formed in areas with little input of clastic material (Papp & Steininger 1974). Sand and carbonate deposition continued throughout the Sarmatian (12.7 – 11.6 Ma) in most parts of the Vienna Basin (Harzhauser & Piller 2002). The general tectonic regime of SW-NE extension continued from the Badenian to the Sarmatian (Figure 2-1 C). The Pannonian (11.6 – 6.2 Ma) period began with a transgression (Harzhauser et al. 2003) covering most of the Sarmatian deposits. Primarily clay and sand was deposited in a lacustrine environment during the Early and Middle Pannonian. During the Late Pannonian the Vienna Basin was filled with alluvial sediments (Strauss et al. 2006)( Figure 2-1 D).

E-W compression and basin inversion (Upper Miocene-Pliocene) In the Latest Pannonian and Pliocene the large-scale stress field with NW-SE orientation shows low-strain N-S shortening for the central part of the Vienna Basin. Instead of N-S compression and E-W extension, an E-W compressive stress field evolved, resulting in basin inversion and sediment deformation (Peresson & Decker 1997).

SW-NE extension (Pleistocene-Recent) Fault-controlled subsidence due to a sinistral trans-tensional regime accompanied by recent seismic activity was detected along the eastern limit of the Vienna Basin (Hinsch & Decker 2003). Faults along the Leitha Mountains (SE of the Basin) are still active today.

2.2 SEDIMENTOLOGICAL PROCESSES The sedimentation in the Vienna Basin started in the Carpathian (17.2 - 16.3 Ma) with sediment input coming from the south. Two major sediment regimes were present during the Carpathian. A marine environment controlled the north (Laa Formation) and a reduced marine to fluviatile sediment regime controlled the south (Aderklaa Formation). After the Carpathian sedimentation phase, an inversion took place and large amounts of the Early Miocene sediments were eroded, in some areas even down to the basement of the basin (e.g. in the Matzen ridge down to the Rhenodanubian Flysch)(Ebner 1997, Steininger & Wessely 2000). A new sequence started with a transgression in the Badenian during which the basin was formed into its present shape. The base of the Badenian sediments upon the Carpathian/Badenian sequence boundary is formed south of the Matzen ridge by the Aderklaa conglomerate representing a lowstand. According to Weissenbäck (1996) the Badenian sediments contain two sequences with several high-stand, transgressive and lowstand systems tracts. The sediment thickness differs largely in depressions or elevations in the basin due to the difference in subsidence rates. The subsidence trend continued throughout the Sarmatian and Pannonian times over the complete Vienna Basin. The first known drainage systems of the Alps discharging into the Vienna Basin are the Aderklaa and the Rothneusiedel conglomerate with gravel (Selge 2005). Discharge started in the northern Vienna Basin from the Middle Badenian with fluviatile transported terrigeneous material forming a submarine fandelta with prominent lobes. A second drainage system originated from the Badenian and is best described as a pre-Danube system. Another river discharge coming from the Carpathian mountain ranges entered the North of the Vienna Basin. Two deltas were entering the basin from the south at the same time. All deltas had their maximum spread in Sarmatian time and can be traced into the Pannonian (Steininger & Wessely 2000). Figure 2-2 shows the 4 different deltaic discharges in a paleogeographic map during the Middle Badenian. Due to the strong tectonic influence and the rapid change between transgression and regression cycles a complex facies differentiation within the Vienna Basin was triggered. Facies diversity was largest in Badenian times. A shallow and a deep marine deposition regime control the Vienna Basin. In the shallow marine regime clastic sequences of conglomerates were deposited through fluviatile transport. In the central part of the basin deep marine sediments were deposited which were arranged in differentiated facies belts from deep water to nearshore sediments. At the rim of the basin and on the elevations within it, coastal terraces are preserved and coral reefs were able to grow. The deeper water sediments consisted of marly shales wherein sand lenses can be found which had been transported from shallower marine locations. In the Sarmatian the deltas and the development of stacked clastic sequences continued. The depositional environment was now more brachyhaline. The Sarmatian shows a more delta-front to delta-slope deposition facies. The subsidence slowed down which helped the prograding deltas to fill the basin with sediments leaving a shallow lacustrine environment during the Pannonian and eventually the final infill of the basin in the Pontian (6.2 – 5.3 Ma). The Pannonian shows a delta-plain to delta-front depositional environment. This fast filling up of the basin is explained due to the fact that the basin is so closely located to the source area of the clastic sediments and the decrease in subsidence of the basin (Piller et al., 2007).

Figure 2-2: Paleogeographical map of the Vienna Basin during the Middle Badenian (15 Ma) after Seifert (1996). The red dot (scale exaggerated) marks the location of the Spannberg 21 well.

3. PALEOMAGNETIC FUNDAMENTALS

3.1 BASIC DEFENITIONS IN PALEOMAGNETISM The magnetic moment per unit volume, M, can be defined by referring either to a pair of magnetic charges or to a loop of electrical current. The magnetic field, H, in a region is defined as the force experienced by a unit positive magnetic charge placed in that region. Both the magnetic moment and magnetic field are expressed in Ampere per meter [A/m]. The magnetic intensity, or magnetization, J, of a material is the net magnetic dipole moment per unit volume. Now we can relate these terms with eachother using the following equation:

푀푖 푖 = 푱 (3.1) 푣표푙푢푚푒

In this equation Mi is the constituent magnetic moment. The total in-situ magnetization of rocks is the vector sum of two components:

푱 = 푱풊 + 푱풓 (3.2) where Ji is the induced magnetization and Jr is the natural remanent magnetism. In paleomagnetism we use the induced magnetization Ji when a material is exposed to the local geomagnetic field H. These quantities are related through the magnetic susceptibility, χ:

푱풊 = 휒푯 (3.3)

The magnetic susceptibility is the net susceptibility resulting from contributions of all minerals but usually dominated by the ferromagnetic minerals (Ch. 3.3) and can be regarded as the magnetizability of a substance. The magnetic susceptibility per unit volume is dimensionless. The vector sum of the magnetic field and the induced magnetization is called the magnetic induction B and is defined as:

푩 = µ표(푯 + 푱) (3.4) in which µo is the permeability of free space in [H/m]. Natural remanent magnetization (NRM) is remanent magnetization present in a rock sample prior to laboratory treatment. NRM depends on the geomagnetic field and geological processes during rock formation and during the history of the rock (Butler, 2004). NRM will be further explained in Ch. 3.5.

3.2 THE GEOMAGNETIC FIELD The Earths‟ magnetic field is best described by two magnetic poles forming a dipole field around the surface of the Earth. The positions of these two poles do not completely coincide with the geographical north and south pole. Instead the geomagnetic poles deviate about 11.5 o from the Earth‟s axis with this angle continuously varying (e.g. Tauxe 2008). The positions of the poles change every day, at present they shift about 90 meters per day (Butler 2004). Due to the changes in position the strength and direction of the field deviates as well. There is also a big change in field strength found between different places on the surface (Tauxe 2008). This is due to the different spreading of isomagnetic lines. Isomagnetic lines are lines which represent equal field strength. In between the poles, approximately at the same latitude as the equator, the spreading is highest, hence the strength of the field is smallest. At the poles, where all isomagnetic lines come together or spread out, the strength is highest. Nowadays the strength is about 6.0*10-5 T near the poles and 3.1*10-5 T near the equator.

Not only can the strength and the direction of the geomagnetic field be changed, it can also be reversed, that is the geomagnetic north pole and south pole change polarity which is called a reversal. Up to now it is still not completely proved and understood what mechanism drives this. The duration of a (sub-chron) reversal is in the order of 104-108 years (Butler 2004).

3.3 MAGNETISM IN ROCKS On the way how a magnetizable particle behaves in a magnetic field we can distinguish three different types of magnetization J, acquired in response to the application of a magnetic field.

3.3.1 DIAMAGNETISM The diamagnetic response is small and opposite (anti-parallel) to an applied field H. The magnetization depends linearly to the applied field, and the magnetic susceptibility is negative. All matter has diamagnetic response, but for substances whose atoms possess atomic magnetic moments, diamagnetism is dominated by effects of magnetic fields on the atomic magnetic moments.

3.3.2 PARAMAGNETISM A paramagnetic mineral contains a magnetic moment parallel to an applied magnetic field, H. The magnetization, J that is hereby acquired is linearly dependent on H. There is no interaction between adjacent atomic moments. As for diamagnetic minerals magnetization will be zero when the applied field H is removed. It follows that the magnetic susceptibility is positive for all paramagnetic minerals. Paramagnetism levels off at high magnetic fields (H>100 T at room temperature).

3.3.3 FERROMAGNETISM Ferromagnetic minerals have a magnetic moment, but unlike the paramagnetic minerals adjacent atomic moments interact strongly. This effect of interaction can produce magnetizations in ferromagnetic minerals that can be orders of magnitude larger than for paramagnetic minerals when they are applied to the same magnetizing field. Each ferromagnetic material has its own saturation magnetization, Js. This is the maximum magnetization a material can acquire when temperature and magnetic field are increased. Ferromagnetic minerals are of importance since they have the ability to record the direction of an applied magnetic field. During removal of the magnetic field, magnetization does not return to zero unlike the paramagnetic and diamagnetic minerals. The path of magnetization as a function of an applied field is called a hysteresis loop. Thus the susceptibility cannot simply be described anymore. Exchange coupling is a result of exchange energy which is caused by the difference of interatomic distances in the crystallography of a solid. Situations occur where parallel coupling within layers of atomic magnetic moments evolve but antiparallel coupling between layers. If the layers have equal magnetic moment, opposing layers cancel, with resulting J=0. This type of coupling is antiferromagnetic (Figure 3-1 b). If layers of unequal magnetic moment are antiparallel, the resulting J points in the direction of the dominant layer. Such materials are called ferrimagnetic (Figure 3-1 c), and many of the important „ferromagnetic‟ minerals are, in fact, ferrimagnetic. Finally ferromagnetic coupling (Fig. 3-1 a) is the result of parallel coupling, i.e. the fields in two adjacent layers point into the same direction (Butler 2004). “Ferromagnetism” is used in the general sense to designate exchange-coupled materials.

Figure 3-1: Exchange couplings for (a) ferromagnetic, (b) antiferromagnetic, and (c) ferromagnetic materials. The net magnetization for a ferrimagnetic material is shown at right; the net magnetization of antiferromagnetic material is zero.

3.4 MAGNETIC MINERALOGY For the description of the mineralogy I will only focus on the magnetic minerals that are assumed to be present in the samples. Discussed are the iron-oxides such as magnetite, maghemite and hematite, the iron-oxyhydroxides, such as goethite and ferrihydrite, and the iron-sulfide greigite.

Iron-oxides Of all iron-oxides, magnetite (Fe3O4) and hematite (Fe2O3) are the most well known. In nature Ti4+ often substitutes for iron in the crystal structure forming titano-magnetite. The total amount of substitution in magnetite and hematite is often denoted in a ternary diagram. In general the magnetization is reduced when Ti4+ ions are present in the crystal structure of magnetite and hematite. Fig. 3-2 shows a ternary diagram for most common iron-bearing minerals.

Iron-oxyhydroxides Goethite (αFeOOH) is the most common magnetic phase in iron oxyhydroxides. Goethite is formed as a weathering product of iron-bearing minerals and as a direct precipitate from iron- bearing solutions. It has a wide occurrence but for the scope of this thesis goethite is not important.

Figure 3-2: Ternary diagram for iron-oxides. The dashed lines with arrows indicate the direction of increasing oxidation. The solid lines are solid solution series (after Tauxe, 2008).

Iron-sulfides The main iron-sulfide that is of relevance for paleomagnetism is greigite (Fe3S4) It occurs in reducing environments and tends to oxidize to various iron-oxides (magnetite and hematite) leaving the paramagnetic pyrite as the sulfide component (Tauxe, 2008). According to Vassiliev et al. (2008) greigite can be subdivided into two groups depending on the two different formation mechanisms, magnetosomal and authigenic greigite. Magnetosomal greigite is formed inside bacteria as monocrystalline minerals i.e. a single bacteria made a single greigite crystal. The greigite producing bacteria prefer reduced conditions and are

probably anaerobic sulphate reducers (Vasiliev et al., 2008). Magnetosomal greigite grains are small (20-75 nm) and very common in sediments. It is shown that magnetosomal greigite can survive geological times and that a primary NRM component can be extracted (Vasiliev et al., 2008). Figure 3-1 shows a photomicrograph of magnetotactic bacteria. Authigenic greigite or early diagenetic greigite is formed as intermediate phase in anoxic and reducing environments to form pyrite (FeS2). Authigenic greigite is relatively large (400-1000 nm) and is formed later than the magnetosomes, deeper in the sediment and therefore acquired a later magnetic field (Vasiliev et al., 2008).

The samples I have looked at were all from a sedimentary origin. The main source for sediments are igneous rocks being transported and weathered for long periods of time. However biological and low-temperature diagenetic agents also work to modify the igneous rocks and they have a significant effect on the magnetic mineralogy in the sediments. It follows that the magnetic components found in sediments may have a large variety in provenance (Tauxe, 2008). (Titano)magnetite may alter when it enters different pH and/or redox conditions. Also, although the geochemistry of seawater is generally oxidizing with respect to the stability field of magnetite, pronounced changes in the redox state of sediments often occur with increasing depth as a function of the breakdown of organic carbon. Such changes may result in locally strongly reducing environments where magnetite may be dissolved and authigenic sulfides are produced such as pyrite (Tauxe, 2008). It was already discussed that the Vienna Basin had periods of high sedimentation rates (up to 1.25 m/kyr in the Upper Sarmatian). High sedimentation rates accompanied by rapid burial of organic matter could lead to a completely anoxic diagenetic environment close to the sediment/water interface (Vasiliev et al., 2008). It is not unlikely that authigenic greigite in the Vienna Basin was formed through bacterial mediation during early diagenesis after deposition. Examples from literature showing the same depositional conditions support this (Vasiliev et al., 2008).

Figure 3-1: Photomicrographs of bacterial magnetites produced by magnetotactic bacteria. a) Intact magnetosome in living bacterium. (Fassbinder et al., 1990) b) Chains recovered from ODP Site 1006D in the Bahamas (Hounslow in Maher and Thompson, 1999)(after Tauxe, 2008).

3.5 NATURAL REMANENT MAGNETISM Luthi (2001) describes three major types of Natural Remanent Magnetism: Thermoremanent magnetism, acquired during cooling in a magnetic field from a high temperature to below the blocking termperature (which is somewhere near but below the Curie temperature). The Curie temperature Tc is the temperature above which ferro- and ferrimagnetic substances behave paramagnetic. When heating a substance atomic distances increase and hence the strength of exchange coupling between atomic magnetic moments decreases with increasing temperature. This reduces the resultant magnetization. When the Curie temperature is reached, interatomic distances are so large that the exchange coupling breaks down. The atomic moments become independent and the material exhibits paramagnetism. Upon cooling below the Curie temperature exchange coupling and ferro- or ferrimagnetism reappear. When a substance is cooled down from above its Curie-temperature to below its Curie-temperature in the presence of

an external magnetic field, the orientation of the newly induced magnetization will be parallel to the orientation of the external field. In the process of thermoremanent magnetization, a bias in the distribution of magnetic moments at higher temperatures freezes when the material drops below the critical temperature, producing remanent magnetization. Chemical remanent magnetism, in which a material, at a temperature below the blocking temperature of thermoremanent magnetism, is chemically altered such that a ferromagnetic mineral is produced. Alternatively, a ferromagnetic mineral might be precipitated from the pore water. Detrital remanent magnetism, which is acquired during deposition, generally through partial alignment of ferromagnetic grains with the earth‟s magnetic field. Sedimentary particles, for example, can be aligned this way while they settle in water. In sedimentary rocks this type of magnetism may be affected by post-depositional reorientation, for example through bioturbation or compaction.

Natural remanent magnetism is important in sedimentary rocks because by measuring this quality one hopes to obtain the direction of the earth‟s magnetic field at the time of deposition, or shortly thereafter. The difficulty of measuring the correct NRM does not lie within measuring the signal, which is roughly one to ten million times smaller than the earth‟s magnetic field, but lies within the distortion of the NRM. Factors controlling the distortion of the signal measured downhole during logging operations are the ferromagnetic drill string and drill bit, the composition of the drilling mud, structurally distortion of the layers, deviation of the wells and plate tectonic drifts that may have occurred since deposition of the rocks of interest. Finally a „blind spot‟ may be produced when inductions Ji and Jr project orthogonally onto the earth´s magnetic field induction B due to geological distortions (Luthi, 2001). Although it might seem impossible to get legitimate data out of measuring the NRM, there are four possible applications for paleomagnetic borehole logging: (i) Age dating; (ii) Identification of unconformities; (iii) Chronostratigraphic well-to-well correlation and (iv) Sequence stratigraphy.

4. METHODS

4.1 SAMPLES The samples were taken from the Spannberg 21 appraisal well. A drilling rock sample was taken every 5 meters and was sealed in an airsealed bag still containing the liquid cuttings. After discussion in the lab in Utrecht I decided to use 32 samples, all at different depths. Together with my supervisor we sought for the best reference depths to be measured. In this search we took variations in lithology, paleomagnetic signal, reversal of geomagnetic field and gamma ray values into account (Figure 5-5). After collecting the sample bags I dried ~100 g of each sample for 3 days in an oven. I used bentonite to simulate the real drilling mud. The temperature in the oven was kept constant at ~50 oC, thus not evaporating the connate water in clay minerals. Now the samples were ready for further processing in the lab in Utrecht ( Figure 4-1).

Figure 4-1: Dried samples just taken out of the oven.

4.2 SUSCEPTIBILITY MEASUREMENT Lab Each sample had large fragments that had to be grinded down to <0.5 mm fragments. This was done using a stone (inert) grinding bucket. Every sample was then collected in small glass jars and weighted with an accuracy of 10-5 g on a microbalans “Sartorius”. Each sample contained ~10 g. The susceptibility meter is of the type “Kappabridge KLY-2” (Figure 4-2). Every sample was measured three times and the apparatus was reset to zero after each measurement. All metal objects such as a watch, phone and jewels were removed because of the high sensitivity of the apparatus. All measurements with the susceptibility meter were carried out using range 2, except for depth 1987.5 m which was carried out at range 3 because of its large susceptibility signal.

Theory The previously explained susceptibility, or magnetizability is measured in this apparatus in 1 direction. Using the least-squared method the final susceptibility vector is calculated. The sensitivity deviates from 0.5 * 10-6 to 200,000 * 10-6 SI within 11 ranges. The glass jars are diamagnetic and their influence has to be taken into account as well as the correction for the drilling fluid. When normalizing the samples to [kg] and [m3], the final SI unit used in this report is [m3/kg].

Figure 4-2: Kappabridge KLY-2 susceptibility meter.

4.3 ALTERNATING FIELD DEMAGNETIZATION Lab Before I was able to measure the samples with the IRM and ARM „robot‟, all samples had to be alternating field (AF) demagnetized. When the susceptibility measurements were carried out, I was able to use the same samples for the demagnetization as well, since the susceptibility meter has no influence on the magnetic properties of the mineral content. Plastic 2x2x2 cm containers were used and filled ~ ¾ of the volume. After weighing the samples all containers are filled with a two component epoxy raisin while stirring the mixture. The stirring is very important because idealistic all grains have to be floating in the epoxy raisin, i.e. not touching neighboring grains. The raisin needs 24 hours to fully harden and the weights are measured again. With an arrow the direction of one central axis through the sample is indicated. The samples are now taken to the AF demagnetizer with Oersted meter of the type “Forster 1.107”. After calibrating the demagnetizer for the geomagnetic field (accuracy 0.5 nT) each sample is slid into the demagnetizer and the sample is demagnetized in three directions of which the one direction indicated with the arrow is of importance for later interpretation of the results.

Theory Demagnetization is the process of demagnetizing the minerals. This can be done in several ways, in our case it is done by applying an alternating magnetic field to the sample that is above the coercivity of the minerals and then slowly decreasing this magnetic field to zero. Coercivity is best described as the intensity of a magnetic field necessary to reduce the magnetization in a certain matter. Therefore coercivity can be best seen as a resistance and it is expressed in Oersted or A/m.

4.4 ISOTHERMAL REMANENT MAGNETIZATION (IRM) Lab The cubic boxes are glued into a second plastic (diamagnetic) container with two open opposite sides. These second containers fit exactly in the ARM/IRM „robot‟: 2G Enterprises DC-SQUID magnetometer (Fig. 4-3). It is called „robot‟ since this apparatus does all the measurements for each sample one by one once the operator has given the order to do so. The samples are placed in a matrix scheme and the robot picks out each sample for measuring the ARM and subsequently the IRM. Once this is done, it places the sample back in the matrix and picks out the following one. The output is the magnetic moment in [A*m2] for a three axis coordinate system. For our data interpretation only the one axis indicated with the arrow will be of importance. Each sample is measured at 60 different magnetic field strengths varying from 0-700 mT.

Theory IRM is the contrary of demagnetization, hence it is an artificial method of creating a magnetization in a mineral. This is done by placing the mineral in a direct magnetic field. When progressively increasing this field and measuring the IRM in the mineral one is able to determine the magnetizable content of the sample. Each mineral has its own characteristic IRM curve when the IRM data is plotted against the applied field. It follows that a mineral must have the capacity to be magnetized which is expressed, as explained in Ch.3.5, as the natural remanence. When using IRM and ARM results I have chosen to normalize with mass and not with volume, since it was easier to weigh the samples than to determine their volume. Hence the unit to be further used for IRM and ARM magnetization in this report is: A*m2 / kg = [Am2/kg].

Figure 4-3: AF Demagnetizer (left) and 2G Enterprises DC-SQUID magnetometer (right)

4.5 ANHYSTERIC REMANENT MAGNETIZATION (ARM) Lab The same „robot‟ that measures the IRM values was also used to measure ARM values.

Theory ARM defers in that way from IRM that it is a measurement for the remanence of a mineral when it is subjected simultaneously to an alternating and a direct magnetic field. ARMs can be produced in two ways during AF demagnetization: - the Earth‟s magnetic field is not adequately cancelled out by the magnetic shielding of the AF demagnetizer - the alternating current passed through the demagnetizer coil does not have a pure wave form. In both cases the resulting small magnetic field will introduce an ARM in the mineral producing a noisy demagnetization path and can obscure the natural remanence. This is especially seen at high demagnetization fields of >50 mT. On the contrary ARMs can also be deliberately produced to investigate the variation in magnetic mineral grain size within a sample. This is done by placing the sample in a constant direct field while an alternating field is applied.

4.6 CURIE TEMPERATURE LOOPS Lab The last measurements that were executed are the Curie temperature loops. Because of the duration of this experiment it was chosen to only take seven samples to be measured. These seven samples were critically picked out, taking care of different lithologies and characteristic magnetic mineral properties. For the measurement only approximately 50 mg of each sample is needed and is carefully embedded in quartz fiber wool. This embedded sample is compacted in a glass tube that is placed in the Curie Balance, which is an inhouse developed instrument. The endpoints and starting points for heating and cooling were given in by the operator. The maximum temperature was set at 700 oC. Each measurement cycle takes up three hours and every four seconds a temperature measurement is taken as well as a magnetic remanence measurement.

Theory By varying the temperature in a vast amount of sample the variation of the magnetization can be measured. By carefully controlling the cooling down curves known magnetic minerals can be detected. Each mineral has its own characteristic Curie temperature. In our measurements Curie temperatures for greigite (340 oC), magnetite (580 oC) and hematite (680 oC) were most relevant. A Temperature-Magnetization- graph for each sample is constructed by the computer making use of Fourier transformations to suppress noise in the signal.

5. RESULTS

5.1 SUSCEPTIBILITY MEASUREMENT Susceptibility measurements were conducted three times on each sample. Sample 1987.5 m had such a high susceptibility that it had to be measured in a higher range (Ch 4.2). As previously stated, benthonite was used to estimate the drilling fluid. To see what the effect is of this matter on the susceptibility, two correction factors are modeled, one which uses a 10 wh% of benthonite and the other one 50 wh%. For sample nr. 15 (1930 m) the susceptibility was lowest, hence benthonite influence was largest: 6% contribution of the total susceptibility in the 10 wh% estimate and 30% contribution of the total susceptibility in the 50 wh% estimate. All other samples give smaller contribution values for benthonite. In Figure 5-1 data of the susceptibility versus depth are plotted.

Susceptibility (10-9 m3/kg) 0 50 100 150 0

200

400

600 10wh% Benthonite 800 50wh% Benthonite 1000

Depth (m) Depth 1200

1400

1600

1800

2000

Figure 5-1: Two different susceptibility values for each depth depending on the contribution of benthonite. Note that sample nr 0 (1987.5 m) is not plotted since its value is ~510*10-9 m3/kg.

5.2 CURIE TEMPERATURE MEASUREMENT The results from the Curie temperature measurements for each sample were plotted on Temperature-Total Magnetization diagrams. These diagrams didn‟t require any manual processing before interpreting them. Because of similarities in the graphs I have chosen to group the seven samples in three groups according to their shapes and outcomes. Group A contains samples nr 7 and nr 26 at depths of 1340 m and 840 m respectively (Figure 5-2). Different cooling down-heating up loops can be seen. The first important loop is the one at a temperature of ~340 oC where greigite is altered (Vasiliev et al., 2008). The cooling-down loop lies beneath the overall heating-up curve so the total magnetization has decreased, indicating that greigite contributed to the total magnetization signal. The next loop is that of pyrite at a temperature of 420 oC. Pyrite and greigite will almost always be found together since pyrite is formed in reducing environments from greigite.

Magnetite formed from pyrite

Max magnetite contribution

No sharp decrease of magnetite altering to hematite at 580oC => maghemite is formed first

Greigite oxidizes: lower magnetization reversal curve

Figure 5-2 : Group A. Curie-Temperature graphs from sample nr. 26 at a depth of 840 m and sample nr. 7 at a depth of 1340 m. Clearly visible is the greigite content at a temperature of 320 oC and the pyrite content at a temperature of 440 oC. Dotted line is the cooling-down (reversal) curve, all other curves are heating-up curves.

Due to the high temperatures pyrite is oxidized to form magnetite (Dekkers pers. communication). One can see from the graph that the signal increases due to the magnetite being formed. The forming of magnetite reaches a maximum at a temperature of 540 oC. At magnetite‟s Curie temperature (580 oC) we expect it to become paramagnetic. When the magnetite particles are large enough they will be oxidized to form hematite. However when the particles are smaller than approximately 35 nm first an intermediate maghemite (γFe2O3) is formed. With increasing temperature maghemite is mineralogically altered to hematite (αFe2O3):

1 2 퐹푒 푂 + 푂 = 3 훾퐹푒 푂 3 4 2 2 2 3

훾퐹푒2푂3 => 훼퐹푒2푂3 (5.1)

Since maghemite has a Curie temperature of 640 oC, the decrease of the signal due to the decrease of magnetite will be spread out over the region 580 – 640 oC. Ultimately at a temperature of 640 oC only hematite is left in the sample and when cooling down one would expect a reversible process, hence a reversible curve that should follow approximately the heating up curve. The cooling down curve (dotted curve in Fig. 5-2) however shows an enormous increase in total magnetization for low temperatures. M. Dekkers has an „ad hoc‟ explanation for this feature: Depending on the size of the hematite grains it will be superparamagnetic. Normally hematite shows an antiferromagnetic character when the grains have sizes varying between 30 – 40 nm. However when the grains are smaller than 30 nm, the grains will behave superparamagnetic. The particles are so small that they will easily be lined up parallel to the applied field and thus contributing to the total magnetization signal. At depth 840 m there is more pyrite present in the sample than at depth 1340 m hence the reversible curve will increase rapidly for small temperatures. Consequently group A consists mainly of greigite and pyrite.

Group B consists of sample nr‟s 9 and 12 at depths of 530 m and 1070 m respectively (Figure 5-3). The first three reversal curves lie beneath the heating-up curve. At 540 oC a very small increase in the reversed curve is visible, which may be a very small content of pyrite. This is visible in both graphs. Typically for these graphs is that the last cooling-down curve lies beneath the heating-up curve and shows a large increase for small temperatures. This can be explained due to small hematite particles formed due to the easy oxidation of small magnetite particles. Consequently group B mainly consists of magnetite and might contain a small amount of pyrite.

Group C contains samples nr‟s 19, 13 and 4 at depths 545, 590 and 1050 m respectively (Figure 5-4). These graphs are similar since they show no greigite or pyrite content. Graph 1050 m is a bit wobbly which could be explained by the fact that the sample was not correctly placed in the measuring apparatus. Graph 590 m shows the bent for the magnetite Curie-temperature at 560 oC, which is slightly low for magnetite (see red arrow in Figure 5-4). It can be explained by a small titanite content in the magnetite crystal structure (Dekkers, per. communication). Graph 545 m shows the magnetite decrease bent exactly at 580 oC. Therefore group C consists mostly of magnetite.

Last cooling-down curve lies beneath heating-up curve => rel. large amount of superparamagnetic hematite => rel. small magnetite particles

Large increase in magnetization => superparamagnetic hematite

Figure 5-3: Group B. Curie-Temperature graphs from sample nr 9 at a depth of 530 m and sample nr 12 at a depth of 1070 m. No greigite content visible; no pyrite visible; magnetite visible at 570-580 oC which oxidises to hematite forming the steep cooling down curve (dotted line).

Pyrite?

Figure 5-4: Group C. Curie-Temperature graphs from sample nr 13 at a depth of 545 m, sample nr 19 at a depth of 590 m and sample nr 4 at a depth of 1050 m. No greigite content visible; no pyrite visible; magnetite visible at 570-580 oC which oxidises to hematite forming the steep cooling down curve (dotted line).

5.3 IRM COMPONENT ANALYSIS IRM measuring results had to be processed into usable data. According to Kruiver et al. (2001) if no magnetic interactions occur, an assemblage of grains of a single magnetic mineral can be characterized by (Figure 5-1): (i) Saturation Isothermal Remanent Magnetization (SIRM), which is the plateau IRM value (ii) B1/2: the field at which half of the SIRM is reached (iii) The dispersion parameter DP: the width of the distribution, given by one standard deviation of the logarithmic distribution (z = 1). If more than one magnetic mineral is present, their IRM acquisition curves add linearly to yield a cumulative curve, provided magnetic grain interaction is negligible (Kruiver et al., 2001). The plots were constructed using the IRM-CLG10 program running in an Excel-workbook. This program is based on the cumulative log-Gaussian analysis (CLG) and constructs a best fit through three different curves: Linear Acquisition Plot (LAP), Gradient Acquisition Plot (GAP) and Standardized Acquisition Plot (SAP). The LAP shows cumulative distribution curves, the GAP is mainly a scatter plot and can be used to approximate the log B1/2 value and the SAP shows the z-score-log B1/2 plot for the different components. The abscissa for all graphs are logarithmic showing log B1/2 values.

Figure 5-5: (a) An example of an IRM acquisition curve, called the linear acquisition plot (LAP). (b) The gradient of acquisition plot (GAP). The dispersion parameter (DP) represents one standard deviation (z=1). (c) The IRM acquisition curve on a probability scale (right-hand ordinate) and corresponding z-score scale (left-hand ordinate), called the standardized acquisition plot (SAP). Solid line (including DP and B1/2) is for the IRM acquisition curve shown in panel (a). Squares are measured data from a single magnetic mineral sample (titanomagnetite). Note that the abscissa is logarithmic for all three plots (After Kruiver et al., 2001).

Plotting of the data starts with the initial values for SIRM, log B1/2 and DP being estimated from the LAP, GAP and SAP. These values are entered in the IRM-CLG 10 program and for the specified distributions the theoretical IRM curves are calculated and added. The modeled LAP, GAP and SAP are compared to the data. The goodness of fit is expressed by the sum of the squared residuals between the data and the model for each plot. The values for SIRM, log B1/2 and DP are optimized interactively by minimizing these squared residuals (Kruiver et al. 2001). The plots contain information of all the remanent magnetizable minerals. In most samples it was not expected to find just one type of mineral, but a mixture of for example magnetite, greigite and hematite. This is of importance before starting the best fit plotting. In the plotting procedure the processor can add as many different components as he wants, each component accounting for one specific mineral. Most plots can either be plotted using two or three components to get a best fit within an error of 1% for each curve. The main component will always account for the largest SIRM signal, however the smaller components must be taken into account as well. In the two component interpretation, the smaller component can be interpreted having the same log B1/2 values but having a large DP (>0.45 mT). The three component analysis though yield a small log B1/2 for the second component and a large log B1/2 for the third component. DP values for both components are neither large nor small (~0.25-0.4 mT). According to Dekkers (pers. communication) it is not proven which method is better so it is assumed both are correct. For this interpretation initially the three component interpretation was used. However some samples showed from the start that they could perfectly be fit using the two component method. This was done for samples nr. 4 (1050 m), nr. 17 (480 m) and nr. 11 (1830 m).

Furthermore SIRM values were scaled to their weight, log B1/2 values transformed to B1/2 values and SIRM for each component was calculated as the percentage of the total SIRM signal (See Appendix for results).

Interpretation of the data Three major minerals were of importance when interpreting the IRM measurements namely magnetite, hematite and greigite. In Table 5.1 typical B1/2 and DP values for these minerals are listed. SIRM is not listed since it is not a mineral specific property. It is a quantitative value for the remanent magnetization. The range in DP values for greigite follows from the different depositional environments. Since magnetosomal greigite is formed within single bacteria, the spreading (DP) will be much smaller since these bacteria have an approximately equal size. Hence the greigite crystals will be (overall) equally shaped and sized as well (Vasiliev et al., 2008).

Mineral B1/2 (mT) DP (mT) Magnetite 25 – 60 ~ 0.35 Greigite (magnetosomal) 70 – 80 ~ 0.20 Greigite (authigenic) 60 – 70 0.25 – 0.35 Hematite 300-800 ~>0.4 Table 5.1: Typical B1/2 and DP values for magnetite, hematite and greigite (Vasiliev et al., 2008, Dekkers pers. communication).

Using the typical values from Table 5.1 and the graphs from the Curie-temperature measurements it should be possible to conclude which minerals were present in the samples. It appeared to be necessary to take the constructed IRM graphs into account as well. Certain IRM graphs got a better fit after a reconstruction using the CLG method. This was done for graphs 530, 1070, 1355 and 1670 m. The Curie-temperature measurements showed that 1070 m did not contain any greigite, however this was estimated from the IRM graph earlier (B1/2 = 79.4, DP=0.16). Taking a second look at the GAP plot revealed a better construction using two components instead of three. This was also done for the other three graphs. Even after this reconstruction of certain graphs, it turned out that several samples had values that could either be greigite or magnetite e.g. sample nr. 24 (1465 m) and nr. 25 (1505 m). Their B1/2 and DP values could be used to identify both greigite and magnetite. Here I have used the DP values to choose for greigite. The DP values suite greigite better than magnetite.

5,0E+06 1,0E+07 IRM_0 IRM_1

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 1,0E+07 1,0E+07 IRM_2 IRM_3

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 1,0E+07 1,0E+07 IRM_4 IRM_5

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 5,0E+06 IRM_6 IRM_7

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 1,0E+07 1,0E+07 IRM_8 IRM_9

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 2,0E+06 2,0E+07 IRM_10 IRM_12

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 2,0E+06 IRM_13 IRM_14

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 2,0E+06 5,0E+06 IRM_16 IRM_17

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 2,0E+04 IRM_18 IRM_34

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3

2,0E+07 5,0E+06 IRM_19 IRM_20

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 2,0E+06 5,0E+07 IRM_21 IRM_22

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 5,0E+06 IRM_23 IRM_24

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 5,0E+07 IRM_25 IRM_26

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 2,0E+07 5,0E+06 IRM_27 IRM_28

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+07 5,0E+06 IRM_29 IRM_30

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3 5,0E+06 2,0E+04 IRM_31 IRM_33

0,0E+00 0,0E+00 0 1 2 3 0 1 2 3

3,0E+06 Raw data

] ] IRM_32 2 2,5E+06

Am Component 1 12

- 2,0E+06 1,5E+06 Component 2 1,0E+06 5,0E+05 Component 3

0,0E+00 synthetic [10 IRM synthetic 0 0,5 1 1,5 2 2,5 3 Sum of 10 components Log B1/2 [mT] Figure 5 -6: IRM acquisition graphs (LAP) constructed using a two component or three component construction method. Graphs are all

Linear Acquisition Plots. Vertical axis show raw IRM values26 [10-12 Am2], horizontal axis show log B1/2 values in [mT]

5.4 CORRELATION OF MAGNETIZATION Three results give an insight in the quantitative magnetic mineralogy which constitutes to the total magnetization: the susceptibility and the SIRM of the IRM and ARM measurement. When these values are plotted against depth, they show approximately the same behavior (Figure 5-6). To compare these results with those seen in many subject related papers, all values have been multiplied with an estimated density of 2200 kg/m3 hence SIRM is now given in [A/m]. Overall the graphs follow the same paths except for the susceptibility graph in the 1550-1930m region. In this region the graph does not correlate with the other two graphs. When this region is compared to the GHMT-log results (left), it follows that the signal should be very small. Therefore the IRM- and ARM- SIRM curves combine best with the GHMT-log. The GHMT-log plotted here shows all the measured data points and is the same as the red line in (Fig. 5.5).

B(1/2) 1st total SIRM/kg Depth (m) Lithology Main component(s) component DP SIRM/total (*10-9 Am2 /kg) (mT) (mT) SIRM (%)

480 siltstone magnetite 39,8 0,36 94 1,20 530 sst, very silty magnetite 31,6 0,38 44 2,47 545 sst, very silty magnetite 60,3 0,28 59 0,73 590 claystone magnetite 39,8 0,37 96 2,43 615 siltstone greigite (M) 75,9 0,22 78 0,29 790 siltstone greigite (M) 70,8 0,24 67 0,57 815 siltstone greigite (M) 79,4 0,18 76 0,60 840 siltstone greigite (M) 75,9 0,19 86 7,49 865 siltstone greigite (M) 79,4 0,19 74 1,72 1035 sst, silty greigite (M) 72,4 0,25 83 0,51 1050 sst magnetite 47,9 0,37 59 2,11 1070 siltstone magnetite 33,1 0,35 45 2,94 1130 siltstone greigite (M) 77,6 0,19 67 0,98 1155 silty claystone greigite (M) 74,1 0,19 78 1,29 1240 siltstone greigite (M) 70,8 0,17 81 3,87 1250 Siltstone greigite (M) 70,8 0,17 93 4,77 1285 siltstone, very sandy greigite (M) 83,2 0,17 64 2,36 1325 sandy clay magnetite 56,2 0,31 85 0,43 1340 sandy clay greigite (M) 70,8 0,21 53 1,66 1355 siltstone magnetite & greigite (A) 39,8/67,6 0,24 49/42 3,52 1440 siltstone greigite (A) 70,8 0,26 42 1,39 1465 siltstone greigite (A) 67,6 0,21 60 1,20 1505 siltstone greigite (A) 61,7 0,26 70 0,65 1640 sst magnetite 56,2 0,28 79 0,38 1670 sst/siltstone magnetite & greigite (A) 31,8/89,1 0,24 46/46 0,39 1820 siltstone/calc sst magnetite 50,1 0,35 53 0,76 1830 siltstone/shaly sst greigite (A) 67,6 0,30 81 0,20 1855 siltstone/shaly sst magnetite 50,1 0,35 79 0,52 1930 calc. sst greigite (A) 64,6 0,30 70 0,28 1987,5 calc sst magnetite 36,3 0,35 90 1,99

Table 5.2: Table with depths, lithology, B1/2 of the first component, DP, SIRM of the first component over the total SIRM value, SIRM per kg and interpreted overall mineral content for each depth sample. Greigite (A) stands for authigenic greigite; greigite (M) stands for magnetosomal greigite. Samples highlighted in blue were measured using Curie temperature measurement.

Table 5.2 shows magnetosomal greigite (greigite M) in the 615 – 1340 m region and alternating magnetite and authigenic greigite (greigite A) mainly in the 1355 – 1930 m region. In the Spannberg 21 well the disconformity separating the Upper Badenian from the Lower Sarmatian is interpreted to be in the 1290-1350m region. Kovac (2004) describes a relative sea-level fall during the Late Badenian to Early Sarmatian. Furthermore Kreutzer (1992) recognizes a low systems tract in the Matzen field during the Upper Badenian and the Lower Sarmatian. A relative sea-level fall may lead to oxic conditions at the lake floor/lake water interface and thus a decrease in authigenic greigite. On the other hand oxic conditions are favoured by the magnetotactic bacteria, hence an increase in magnetosomal greigite could be expected. Indeed this is the case from ~1350 m and up into the Sarmatian. Further examining the results from table 5.2 shows no direct correlation between lithology and the most dominant magnetic mineral. Siltstone is most abundant in the lithology and both magnetite and greigite are found in this type of sediment. The sandstones (sst) seem to contain mainly magnetite whereas other types of lithology are so scarce that it is impossible to interpret them with any validity. Figure 5-5 shows an extended log with combined GHMT-log, lithology and interpreted magnetic mineral content.

Figure 5-5: Log showing from left to right: interpreted dominant magnetic mineral, Gamma Ray, litholog, depth, GHMT susceptibility and GHMT remanent magnetism.

-6 -6 Filtered GHMT Susceptibility (*10 SI) ARM SIRM ( A/m) IRM SIRM ( A/m) Susceptibility (*10 SI) 0 200 400 0 50 100 0 50000 100000 0 10.000 20.000

0 Depth(m)

200

400

600

800

1000

1200

1400

1600

1800

2000 Figure 5-6: Log results GHMT (left)(bleu line is measured data, red line contains only the sample points as used in the following graphs), SIRM ARM (middle-left), SIRM IRM (middle-right) and KLY-2 susceptibility (right). Note that the value for sample nr 0 (1987.5 m) is not plotted (was not measured in GHMT log). Note that susceptibility has units [SI] and SIRM is given in [A/m] using an estimated density of 2200 kg/m3.

6. DISCUSSION & CONCLUSION

The sediments deposited during the Badenian, Sarmatian and Pannonian in the Spannberg 21 well reflect a variety in magnetic properties and lithologies. When using four different measuring techniques to measure different magnetic mineral properties, the identification of the main magnetic mineral content of each sample succeeded. An explanation is found for the question why the GHMT-tool measured so little remanent magnetism at certain depths.

IRM acquisition curves indicated the presence of mainly three different components. In most of the cases one or two components were fitted to account for the skewness on the left- and/or right- hand side of the acquisition curves. Depending on the usage of either the two or three component analysis, one or two components were interpreted as either greigite, magnetite or both present in a sample. The measuring apparatus is a very accurate tool. It is not likely to find any significant errors in this tool. Sometimes the values for B1/2 and DP were overlapping typical values for both magnetite and greigite. Therefore it was necessary to perform an additional rock-magnetic experiment.

The Curie-temperature measurements were carried out up to temperatures of 700 oC. However at temperatures over 400 oC the magnetic mineralogy starts to get affected. According to Deng et al. (2004) formation of magnetite from iron-containing silicates/clays or by reduction due to the burning of organic matter is possible. It is clearly visible that the sample material oxidizes, since it returns out of the apparatus bearing a red color, while it goes in with a different color. Normally one would expect a more or less perfect reversible process, but the contrary is true. Due to the formation of superparamagnetic hematite, the cooling-down curve will not onlap the heating-up curve. Five samples out of seven match with the interpretations from the IRM curves. Samples 530 m and 1070 m were interpreted as greigite from the IRM curves, but the Curie temperature results showed detritic magnetite and hardly any presence of greigite. The IRM curves for these and two other depths, 1355 m and 1670 m, were reconstructed and a better matching result was found. The mismatch has several explanations. First the construction of the IRM graphs is very subjective. It all depends on the „eye‟ of the interpreter. Since one is able to use two or three (or even more) components to fit the LAP, GAP and SAP graphs, there is no theory that states which is best. In 12% of the samples (=4 samples) I used a two component model, whereas in the other models I used three components. Of all four reconstructed graphs, two were reconstructed using a two component fit instead of three. The sum of squared residuals calculated for the two or three component fitting was equal for each fitting method. Secondly the starting state of the magnetic system, magnetic interaction and thermal relaxation have a strong influence on the form of the IRM curve, meaning that in a substantial number of cases the lognormal assumption fails and the CLG procedure can produce misleading interpretations (Heslop et al., 2004). Furthermore the CLG procedure will provide useful rock-magnetic information when samples contain well separated magnetization distributions, however, strongly overlapping distributions or very minor components must be treated with caution (Heslop et al., 2004). For magnetic minerals with similar coercivities (e.g. magnetite and authigenic greigite) additional rock-magnetic tests could be required (Kruiver et al., 2001). In this study an additional Curie-temperature test was only carried out on seven samples out of 32. Finally the last error which may lead to differences in Curie-temperature results and the CLG- procedure is due to the small sized sample that is used for the Curie-temperature measurement. Only 50 mg (or less) of each sample is used for the experiment. Magnetic minerals are trace constituent. Although a sample of 50 mg still contains over a million grains, the sample is rather small. For the qualitative magnetic mineral analysis one sample gives a first insight, however for a more accurate quantitative analysis of the magnetic minerals the experiment should be executed four times. (Dekkers, pers. communication).

Even when a match was found between the Curie-temperature measurement and the IRM curves, certain samples were overlapping typical B1/2 and DP values for magnetite and greigite. Sample 1505 m (D1/2 = 61.7, DP=0.26) can be interpreted both as magnetite or as greigite using literature values from Table 5.1. Using only the data available in this study, I have interpreted this

sample to be greigite bearing because of the low DP value which suits greigite better than it does magnetite.

A problem that was encountered when running the GHMT log was the low signal at certain depths. Whether this was an instrumental/logging error or not was verified by comparing GHMT-log with IRM, ARM and susceptibility measurement results. Results from these three measurement techniques support the GHMT results. The remanent magnetic signal at certain depths can be very small. It was found that magnetite and greigite are the main constituters of the total remanent magnetization. However, the formation of authigenic greigite occurs at a later stage, deeper in the sediment and therefore can have a different polarity than the already present magnetite and/or magnetosomal greigite. When this is the case, part of the remanent magnetization or even the complete magnetic signal can be cancelled out (Dekkers, pers. communication). Greigite is formed in anoxic conditions, which can occur in deep lake waters or during high sedimentation rates. When sedimentation rates are small, less greigite may be formed. The same holds for the oxygen content of the water. When the water contains sufficient oxygen, no greigite will be formed at the sediment/water interface. The overall result can be a smaller content of greigite and hence a smaller GHMT signal. However this may be compensated by detritic magnetite. When sedimentation rates are high, larger amounts of magnetite will be transported to the basin constituting to a larger magnetization, provided the sediment provenance shows sufficient iron bearing rocks. It appeared not to be possible to link lithology and the most abundant magnetic mineral in the layers.

To get more certainty on the exact magnetic mineral content, it may be good to use other rock- magnetic properties techniques. Electron Microscopy may lead to a more detailed description of the mineral content (Vasiliev et al., 2008). Furthermore X-ray diffraction is a method to detect crystallographic structures and is therefore useful to determine the mineral content. More Curie- temperature measurements could also be done since the validity of extrapolating only one small sample to reservoir scale is questionable.

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OMV (2007) Intent to Drill, Appraisal Well, Spannberg 21, Internal company report.

Figure 4-2: http://www-odp.tamu.edu/publications/tnotes/tn34/tn34_f18.htm#29615

APPENDIX

Sample Depth (nr) (m) 0 1987,5 1 1155 2 865 4 1050 5 1440 6 1355 7 1340 8 1285 9 530 10 Bent 11 1830 12 1070 13 545 14 1640 15 1930 16 1670 17 480 18 1035 19 590 20 1855 21 615 22 1250 23 1325 24 1465 25 1505 26 840 27 1240 28 1130 30 790 31 1820 32 815 33 Blanco 34 Blanco Table A: Table can be used to convert sample nr’s to sample depths.

Table A-2: Measured and calculated IRM values for SIRM, B1/2 and DP for components 1,2 and 3.