Natalia S Bezaeva
email: bezaeva@physics.msu.ru

Jérôme Gattacceca
email: gattacceca@cerege.fr

Pierre Rochette
email: rochette@cerege.fr

Ravil A Sadykov
email: rsadykov@inr.ru

Vladimir I Trukhin
email: trukhin@phys.msu.ru

Demagnetization of terrestrial and extraterrestrial rocks under hydrostatic pressure up to 1.2 GPa

Keywords: Pressure demagnetization, Hydrostatic pressure, Remanent magnetization, Rock magnetism

We carried out hydrostatic pressure demagnetization experiments up to 1.24 GPa on samples of terrestrial and extraterrestrial rocks and minerals of different lithologies as well as on synthetic samples. The magnetic remanence of samples was measured directly under pressure using a non-magnetic high pressure cell of piston-cylinder type that was inserted into a high sensitivity SQUID magnetometer. In order to bring light on the pressure demagnetization effect, we investigated 50 samples with different magnetic mineralogies, remanent coercivities (B cr ) and hysteresis parameters. The samples consisted of pyrrhotite-, magnetite-and titanomagnetite-bearing Martian meteorites, taenite-, tetrataenite and kamacite-bearing ordinary chondrites and pyrrhotite-bearing Rumuruti chondrite; magnetite-and titanomagnetite-bearing basalts, andesites, ignimbrites, obsidians and granites; a variety of pyrrhotite-and hematite-bearing rocks and minerals (jasper, schist, rhyolite, radiolarite); samples of goethite and greigite as well as synthetic samples of dispersed powders of magnetite, hematite, pyrrhotite and native iron set into epoxy resin. Under hydrostatic pressure of 1.24 GPa, applied in a low magnetic field (<5µT), the samples lost up to 84% of their initial saturation isothermal remanent magnetization (SIRM) without any changes in their intrinsic magnetic properties. We found that the efficiency of the pressure demagnetization is not exclusively controlled by the magnetic hardness of the samples (B cr ), but that it is strongly dependent on their magnetic mineralogy. For a given magnetic mineralogy the resistance to hydrostatic pressure is roughly proportional to ln(B cr ). It was shown that there is no simple equivalence between pressure demagnetization and alternating field demagnetization effects.

Introduction

Hypervelocity impacts are a major mechanism for the evolution of the solid matter in our solar system. Shock waves generated during impacts can modify both intrinsic magnetic properties (Gattacceca et al., 2007a;[START_REF] Louzada | Effect of shock on the magnetic properties of pyrrhotite, the Martian crust and meteorites[END_REF][START_REF] Nishioka | Shock-induced anisotropy of magnetic susceptibility: impact experiment on basaltic andesite[END_REF][START_REF] Gilder | Systematic pressure enchancement of titanomagnetite magnetization[END_REF] and remanent magnetization (Pohl, 1975;[START_REF] Kletetschka | Pressure effects on Martian crustal magnetization near large impact basins[END_REF][START_REF] Gattacceca | Investigating impact demagnetization through laser impacts and SQUID microscopy[END_REF][START_REF] Gattacceca | On the efficiency of shock magnetization processes[END_REF][START_REF] Louzada | Effect of shock on the magnetic properties of pyrrhotite, the Martian crust and meteorites[END_REF] of rocks. Consequently, the magnetic record of solid bodies in the solar system, affected by impacts to different degrees, could have been erased or overprinted by shock events. Understanding the process and the physical mechanism of the impact remagnetization is therefore a key issue to the interpretation of the crustal magnetization of Mars [START_REF] Hood | Distribution of crustal magnetic fields on Mars: shock effects of basin-forming impacts[END_REF][START_REF] Hood | Magnetic anomalies near Apollinaris Patera and the Medusae Fossae Formation in Lucus Planum[END_REF], the Moon [START_REF] Cisowski | Magnetic effects of shock and their implications for lunar magnetism (II)[END_REF][START_REF] Halekas | Demagnetization signatures of lunar impact craters[END_REF][START_REF] Halekas | Magnetic fields of lunar multi-ring impact basins[END_REF], small solid solar system bodies such as asteroids [START_REF] Chen | Hypervelocity impacts and magnetization of small bodies in the solar system[END_REF] as well as paleomagnetic records of meteorites and extraterrestrial materials available from sample return. Concerning the Earth, shock-induced changes in rock magnetic properties and magnetic remanence should be taken into consideration while studying the remanent magnetism of terrestrial impacts [START_REF] Halls | The State of Islands meteorite impact site: a study of shock remanent magnetization[END_REF][START_REF] Pesonen | Palaeomagnetism of the Lappajärvi impact structure[END_REF][START_REF] Pilkington | The geophysical signature of terrestrial impact craters[END_REF][START_REF] Louzada | Paleomagnetism of Lonar impact crater, India[END_REF].

Different authors have carried out experimental investigations of shock demagnetization (remagnetization) of rocks and pure minerals in the 1-30 GPa peak pressure range. Different techniques have been used for shock waves generation: air or gas gun accelerating aluminium or copper projectiles [START_REF] Hornemann | A compressed air gun accelerator for shock magnetization and demagnetization experiments up to 20 kbar[END_REF][START_REF] Pohl | Shock magnetization and demagnetization of basalt by transient stress up to 10 kbar[END_REF][START_REF] Martelli | Hypervelocity cratering and impact magnetization of basalt[END_REF][START_REF] Cisowski | The effect of shock on the magnetism of terrestrial rocks[END_REF][START_REF] Srnka | Magnetic field and shock effects and remanent magnetization in a hypervelocity impact experiments[END_REF][START_REF] Dickinson | Shock magnetism in fine particle iron[END_REF][START_REF] Louzada | Effect of shock on the magnetic properties of pyrrhotite, the Martian crust and meteorites[END_REF]; high explosive and nuclear charges [START_REF] Hargraves | Investigations of the effect of shock on natural remanent magnetism[END_REF][START_REF] Pesonen | Magnetic properties of diabase samples shocked experimentally in the 4.5 to 35 GPa range[END_REF] A c c e p t e d M a n u s c r i p t 3 Gattacceca et al., 2007a); free falling mass [START_REF] Kletetschka | Pressure effects on Martian crustal magnetization near large impact basins[END_REF] and pulsed laser [START_REF] Gattacceca | Investigating impact demagnetization through laser impacts and SQUID microscopy[END_REF][START_REF] Gattacceca | On the efficiency of shock magnetization processes[END_REF]. The main caveats of such experiments are the complexity of dynamic pressure calibration, the possible mechanical damages of investigated samples, and deciphering of the effect of deviatoric versus hydrostatic stresses. Indeed, it is known that remanent magnetization is more sensitive to non-hydrostatic (deviatoric) than hydrostatic stresses [START_REF] Nagata | Main characteristics of piezo-magnetization and their qualitative interpretation[END_REF][START_REF] Martin | The influence of stress path on thermoremanent magnetization[END_REF]. Moreover, shock may permanently modify the intrinsic magnetic properties (e.g., coercivity, see Gattacceca et al., 2007a) thus complicating the interpretation.

As for meteorites, considering the relative rarity of extraterrestrial material on the Earth, it is excluded for most of them to perform shock experiments that are not only destructive but also require rather large sample volume. Numerous parameters must be considered when studying the effect of shock on the magnetic remanence: shock intensity and duration, background magnetic field during the shock event, magnetic mineralogy, pre-shock magnetization and temperature. This large number of parameters, that are sometimes difficult to control, complicates the comprehension of shock effect on rock magnetic remanence.

Static pressure experiments are well suited to tackle these problems. They allow better pressure calibration and can be non-destructive for samples. However, they were until recently limited to the low pressure range for under pressure measurements (<0.1 GPa, e.g., [START_REF] Pozzi | Effets de pression en magnétisme des roches[END_REF].

Experiments were also carried out by pressurizing the sample up to 2 GPa, and remeasuring the remanence outside the pressurizing device (Pearce and Karson, 1988). More recently Rochette et al. (2003) compressed a pyrrhotite sample up to 3 GPa in a piston-cylinder press and remeasured isothermal remanent magnetization (IRM) after pressure release. It was found that pyrrhotite undergoes a high pressure magnetic transition under a pressure of 2.8 GPa, which results in a complete loss of its magnetic remanence. This experimental scheme has the disadvantage of needing a new sample and few days of experiments per each pressure value. Moreover, these experiments, by using a solid confining media, generate some deviatoric stress on the sample.

A c c e p t e d M a n u s c r i p t 4 [START_REF] Gilder | Static stress demagnetization of single and multidomain magnetite with implications to meteorite impacts[END_REF] performed IRM measurements of pure single domain (SD) and multidomain (MD) magnetite under quasihydrostatic load up to 4.2 GPa using a diamond anvil non-magnetic cell (in the Earth's magnetic field) and also observed a pressure demagnetization effect. [START_REF] Gilder | Systematic pressure enchancement of titanomagnetite magnetization[END_REF] carried out pressure experiments up to 6 GPa using a moissanite anvil cell on natural and synthesized MD titanomagnetite with different titanium concentration, but this work was focused upon the influence of stress on the acquisition of IRM.

All these experiments are restricted to pure strongly magnetic minerals due to the minute sample size (e.g., for the diamond anvil cell the cylindrical sample chamber was 400 m in diameter and 100 m in height) and cannot be realized on bulk rock samples without extracting their magnetic fraction.

Pressure demagnetization experiments on bulk rock samples have significant implications in solid-state physics and geophysics, in particular in paleomagnetism and interpretation of crustal magnetic anomalies of the solid Solar System bodies. As crustal rocks suffer the load created by overlying rocks and/or water column (for instance 0.06 GPa for 5 km of water and 350 m of sediments), laboratory studies of the effect of pressure on the remanent magnetism of rocks may be helpful for the comprehension and interpretation of the paleomagnetic signal of the deep seated rocks and crustal magnetic anomalies. However, together with pressure, crustal rocks undergo the concomitant influence of high temperatures, making the situation even more complex. At pressures up to 1.5 -2 GPa, which corresponds to a crustal thickness of 50-70 km, titanomagnetites do not crystallize any more [START_REF] Valeev | Magnetic properties of rocks at high pressures and temperatures and the problem of magnetoactive layer of the Earth's crust and upper mantle[END_REF]: this is the upper limit of relevant pressures.

Despite previous works, the effect of pressure on the remanent magnetization is still poorly known for natural materials for pressures of the order of 1 GPa. The goal of this work is to present a thorough investigation of the effect of hydrostatic pressure up to 1.24 GPa on the magnetic remanence of rocks within a wide range of magnetic mineralogies. We investigated 50 samples of terrestrial and extraterrestrial rocks and minerals as well as synthetic samples with the A c c e p t e d M a n u s c r i p t 5 following magnetic carriers: magnetite, titanomagnetite, hematite, pyrrhotite, native iron and nickel iron, goethite, greigite. For each magnetic mineralogy we studied different samples spanning a wide range of remanent coercivity (B cr ).

Samples and Measuring Techniques

Experimental Setup

In order to isolate the pressure demagnetization effect on rock magnetic remanence from the creation of piezo-remanent magnetization after pressure application (studied in many previous works, e.g., [START_REF] Nagata | Main characteristics of piezo-magnetization and their qualitative interpretation[END_REF][START_REF] Kinoshita | Studies on Piezo-Magnetization (III) PRM and Relating Phenomena[END_REF][START_REF] Pozzi | Effets de pression en magnétisme des roches[END_REF] we always applied the pressure in a low magnetic field (<5 T).

The experimental setup was designed for room temperature measurements of magnetic remanence of relatively large rock samples (up to 5.8 mm in diameter and 15 mm long cylinders) under hydrostatic pressure up to 1.24 GPa. We used a non-magnetic composite high-pressure cell of piston-cylinder type with an inner diameter of 6 mm allowing direct measurement in a SQUID magnetometer (Fig. 1). The cell was made of titanium alloy and "Russian alloy" (Ni 57 Cr 40 Al 3 ).

Samples were placed into a teflon capsule locked with a teflon plug and filled with inert polyethylsiloxane (PES-1) liquid allowing converting the uniaxial pressure on the pistons into a pure hydrostatic pressure [START_REF] Kirichenko | Properties of polyethilsiloxane as a pressure-transmitting medium[END_REF]. Pressure inside the cell was calibrated with a manganin pressure sensor. Details of the cell design can be found in [START_REF] Sadykov | Nonmagnetic high pressure cell for magnetic remanence measurements up to 1.5 GPa in a superconducting quantum interference device magnetometer[END_REF].

After loading of the cell with a press (Graseby Specac 15011), pressure was locked inside the cell. The press with the cell inside was placed at the center of three pairs of perpendicular Helmholtz coils connected to stabilized DC supplies. Due to the presence of mobile metallic parts in the press, it was not possible to obtain a stable very low ambient field. The magnetic field in the area of the investigated sample was monitored using a 3 axis flux-gate magnetometer and was always below 5 T. Measurements of the magnetic moment of the sample under A c c e p t e d M a n u s c r i p t 6 pressure were performed by inserting the pressure cell containing the sample in a 2G Enterprises SQUID magnetometer. This magnetometer allowed measuring the remanent magnetic moment up to 10 -4 Am 2 with a noise level of 10 -11 Am 2 .

We investigated the intensity and stability of the magnetic remanence of the cell itself without and under pressure. The remanent magnetic moment of the cell remained stable at 310 -8 Am 2 at zero pressure and up to 510 -8 Am 2 under pressure in the pressure range 0 to 1.24 GPa, after 5-10 min of relaxation in a magnetically shielded room (ambient field ~100 nT). All measurements of magnetic remanence under pressure presented in this work were performed after 5-10 min of cell relaxation inside the SQUID magnetometer (ambient field 1-2 nT).

In this study the remanence of the investigated samples was always at least one order of magnitude above the remanence of the cell. In most cases, the samples were given a saturation isothermal remanent magnetization (SIRM) in a 3 T magnetic field (9 T for goethite, greigite and hematite-bearing samples) by using a pulse magnetizer MMPM9 from Magnetic Measurements Ltd. before pressure application. A few pressure demagnetization experiments were carried out on a laboratory imparted thermoremanent magnetization (TRM) or a natural remanent magnetization (NRM). The pressure was increased stepwise (8 steps up to 1.24 GPa) and the remanence was measured after each step. The pressure was then released down to zero and the remanence measured again. The specimen was then extracted from the pressure cell and its residual remanence (IRM PR ) was demagnetized by alternating field (AF) up to 150 mT and measured using the same cryogenic magnetometer. The sample was then saturated again and was demagnetized by AF. All these measurements were carried out at room temperature.

In some cases the maximum pressure (1.24 GPa) was applied and released several times to check for the effect of repeated loading on the magnetic remanence. In most cases in addition to all the above described measurements, AF demagnetization curves of SIRM before compression were also measured. In our experiments pressure application did not produce any mechanical destruction of the investigated samples.

A c c e p t e d M a n u s c r i p t 7 Before and after pressure application the investigated samples were subjected to standard magnetic analyses. Measurements of low field magnetic susceptibility  0 and thermomagnetic curves up to 700 C were performed with KLY2-CS2 and MFK1-CS3 Agico apparatus (we used different subsamples for heating experiments). Hysteresis loops at room temperature and remanence measurements were obtained with a Micromag Vibrating Sample Magnetometer (VSM) with maximum applicable magnetic field 1 T. Curie temperatures (T C ) of the investigated samples were obtained from thermo magnetic curves. For titanomagnetite-bearing samples substitution rate (ulvospinel content x) was estimated using the corresponding T C values according to [START_REF] Hunt | Rock Physics and Phase Relations: A Handbook of Physical Constants[END_REF].

Description of Samples

Pressure demagnetization experiments were carried out on samples of meteorites, terrestrial rocks and minerals as well as synthetic samples with the most common magnetic carriers (titanomagnetite Fe 3-x Ti x O 4 ; magnetite Fe 3 O 4 ; hematite Fe 2 O 3 ; pyrrhotite Fe 1-x S; goethite FeOOH; greigite Fe 3 S 4 ; kamacite Fe 0.9 Ni 0.1 ; taenite Fe 0.7 Ni 0.3 ; tetrataenite Fe 0.5 Ni 0.5 ). The main magnetic properties of the investigated samples are presented in Table 1.

We investigated Martian meteorites - [START_REF] Bezaeva | Pressure demagnetization of the Martian crust: Ground truth from SNC meteorites[END_REF] NWA 998 nakhlite and Los Angeles and NWA 1068 basaltic shergottites, nickel iron-bearing ordinary chondrites -Bensour, Pultusk and Saratov -and a pyrrhotite-bearing Rumuruti chondrite NWA 753.

Regarding terrestrial rocks we investigated mostly igneous rocks: three types of andesites from Sardinia (Italy), au1b from Monte Au, cug1b from Monte Cugguruntis and osb10b from Monte Osilo, and five types of basalts: Pleistocene sample ba from the Bas-Vivarais area (France), Pleistocene alkaline sample bb from Chanteuges (Haute-Loire, France), Oligocene sample be from Ethiopian traps and two unshocked samples (pd6-2-1, pd6-2-4) from Lonar impact crater (Maharashtra state, India). We have also investigated a Pleistocene obsidian (kil-2)

A c c e p t e d M a n u s c r i p t 8 from the Kilimandjaro volcano (Tanzania) as well as four granites from Corsica (France), five samples of ignimbrites (all from Sardinia, iro8 from Monte Ironi and mtd5b from Monte Torru, ona12 and spi3301, spi3903 from San Pietro Island) and a sample of jasper (from San Pietro Island, Sardinia) as well as some samples of radiolarite from the French Alps, rhyolite from Esterel range (France) and Precambrian-Ordovician metamorphic schist from Wilson Terrane (Northern Victoria Land, Antarctica). We have also studied a natural hydrothermal fibrous goethite sample GT from Tarn (France) with only 0.42 wt.% of impurities (including 0.16% H 2 O; 0.12% SiO 2 and Ca, K, Ga for the >100 ppm elements, after Rochette et al. 2005b), and two samples of pyrrhotite from the Harvard Mineralogical Museum (Cambridge, USA): monocrystal 127037 from Chihauhau, Mexico and monocrystal 98080 from Sudbury (Canada).

Two natural samples of SD greigite have also been investigated under pressure. Sample greig-tw1 comes from the Plio-Pleistocene marine sediments from the Lower Gutingkeng Formation in southwestern Taiwan and sample greig-it1 is from the upper Pliocene marine sediments from the Valle Ricca section near Rome, Italy. Synthetic samples were in form of dispersed powders set into epoxy resin. Synthetic samples of monoclinic pyrrhotite had grain size x(m) [75,100] for sample a and x(m) [150,250] for sample b. Synthetic samples of dispersed powders of magnetite with x<25 m, iron with x=10 m and MD hematite were prepared using epoxy resin Araldite AY 103 by VANTICO.

Different mass concentrations (0.3; 0.5; 0.8; 17%, see Table 1) of magnetite powder were used for the preparation of synthetic samples of magnetite in order to check whether the pressure resistance of magnetite is dependent on the magnetic interaction of grains. Mass concentration of synthetic samples of iron and hematite was of 2.5% and 2.1% respectively. For magnetite and iron we assume that clustering is not a serious issue as the grains are multidomain and stiring the viscous loaded resin should be sufficient to destroy magnetic aggregates. While magnetite and iron powders are commercial synthetic products, the hematite powder comes from a natural ore from Elba Island (Italy); the pyrrhotite samples are described in [START_REF] Dekkers | Magnetic properties of natural pyrrhotite part 1: Behavior of initial susceptibility and saturation-magnetization related parameters in a grain-size dependent framework[END_REF]. The typical mass for the investigated rock samples was in the 0.02 to 0.5 g range (Table 1). As one can see from Table 1, we have investigated under pressure numerous samples with different magnetic mineralogies within a wide range of B cr and grain sizes: from single domain, pseudosingle domain (PSD) to multidomain.

Experimental results

Main characteristics of the pressure demagnetization experiments

The used pressure cell is characterised by a relatively low but nonzero remanent magnetic moment (see above). In order to check the need for correction of the magnetic remanence of investigated samples by the cell magnetic remanence, the sample of rhyolite (rb7a) was chosen.

This sample is the better suited as it was shown to be the least magnetic (initial SIRM inside pressure cell is 5.4810 -7 Am 2 ). Fig. 2 displays the evolution of the magnetic moment of the cell versus pressure (after 5, 10 and 15 min of the cell relaxation in the magnetically shielded room) as well as an uncorrected and vectorially corrected pressure demagnetization curves for rb7a sample. As one can see from Fig. 2, the correction in case of the least magnetic sample rb7a appears to be negligible. Therefore, pressure demagnetization curves for all other samples (with much higher magnetic remanence, see Table 2) were not corrected for the remanent magnetic moment of the cell.

The reproducibility of our pressure demagnetization experiments was tested in two different ways. Firstly, using exactly the same experimental protocol, we investigated under pressure the magnetic remanence of two subsamples of radiolarite (radiol-1 and radiol-2) and two subsamples of Bensour meteorite (Bensour-i and Bensour-j) originated from the same bulk samples. The pressure demagnetization curves for the two radiolarite subsamples are very close (Fig. 3). The same holds for Bensour subsamples. The observed minor differences may come from sample heterogeneity as well as from the uncertainty on the applied pressure value.

A c c e p t e d M a n u s c r i p t 10 Secondly, we used twice the same synthetic sample of dispersed powder of magnetite set into epoxy resin (sample magn6c, see Table 1) for two successive pressure demagnetization experiments with an identical experimental protocol. The identical pressure demagnetization curves further demonstrate the reproducibility of our experiments.

As mentioned above (see 2.1), we used the eight-step experimental protocol in the pressure range 0 to 1.24 GPa (see Fig. 4a-f). Several experiments were carried out in order to check the possible influence of the chosen experimental protocol (number of pressure steps up to the maximum pressure) on the final pressure demagnetization effect. Figure 5 displays the onestep pressure demagnetization curves of SIRM of microdiorite up to 0.31, 0.62, 0.93 GPa and the eight-step pressure demagnetization curve up to 1.24 GPa. Each curve corresponds to a different microdiorite subsample. As seen from Fig. 5, the pressure demagnetization degree at a given pressure value p is not sensitive to the number of pressure steps up to p. This points to the additivity of partial pressure demagnetization, consistent with previous work (e.g., [START_REF] Pearce | On pressure demagnetization[END_REF].

In order to show that in our experiments the pressure demagnetization effect is not dependent on the duration of pressure application, we carried out the following experiment on a sample of basalt ba (see Table 1). The sample was imparted an SIRM and was then put under 0.93 GPa, which led to 48% decrease in its initial magnetic remanence. The sample was then left under pressure during eight days; no further decrease of its magnetic remanence was observed.

Thus, we conclude that the efficiency of pressure demagnetization is not time-dependent.

Most pressure demagnetization experiments were carried out on magnetically saturated samples (SIRM). It is known that IRM is not equivalent to NRM or laboratory induced TRM in terms of coercivity spectrum [START_REF] Dunlop | Rock Magnetism: Fundamentals and Frontiers[END_REF]. Therefore, these different types of remanence may not have the same pressure sensitivity. NRM is expected to be more pressure resistant than IRM and TRM as it is generally already partially demagnetized by viscous decay.

We carried out pressure demagnetization experiment on IRM, NRM and laboratory induced A c c e p t e d M a n u s c r i p t 11 TRM of basalts (ba and bb). NRM was confirmed to be the most pressure resistant remanence (Fig. 6, sample bb). Our results also showed that in some cases IRM was more easily demagnetized by hydrostatic pressure than TRM (Fig. 6, sample ba), consistent with [START_REF] Avchyan | Effect of hydrostatic pressure up to 8000 kg/cm 2 on various types of remanent magnetization of rocks (translated from Russian)[END_REF]; but this is not a general result (see Fig. 6, sample bb and also [START_REF] Pohl | Shock magnetization and demagnetization of basalt by transient stress up to 10 kbar[END_REF].

Hydrostatic pressure of 1.24 GPa did not produce any changes in SIRM of our samples.

For example, both AF demagnetization curves of initial (pre-compressed) SIRM and postcompression SIRM of granite sample bf9804 are identical (Fig. 7). We do not observe any modification of the intrinsic magnetic properties (hysteresis parameters, low field magnetic susceptibility  0 , B cr ) for the pressure range used in this work, since these modifications occur for pressures above several GPa (Gattacceca et al., 2007a;[START_REF] Louzada | Effect of shock on the magnetic properties of pyrrhotite, the Martian crust and meteorites[END_REF][START_REF] Nishioka | Shock-induced anisotropy of magnetic susceptibility: impact experiment on basaltic andesite[END_REF]. Therefore, the grains are not affected by irreversible changes in the crystalline structure.

Thus, pressure demagnetization is more likely due to domain wall displacements (irreversible upon decompression). If so, domain walls are more likely reset by a new saturation, as for all investigated samples SIRM before compression was found to be identical to SIRM imparted after decompression. As it was shown in [START_REF] Bezaeva | Pressure demagnetization of the Martian crust: Ground truth from SNC meteorites[END_REF] for the Martian meteorites case, pressure sensitivity of rocks is mainly controlled by their magnetic mineralogy. Therefore, demagnetization curves (Fig. 4a-f) were grouped by magnetic mineralogies irrespective of sample type (natural or synthetic, terrestrial or extraterrestrial origin etc.).

Systematic study of the pressure demagnetization effect

A c c e p t e d M a n u s c r i p t 12

The effect of permanent pressure demagnetization (decreasing remanence) was observed on all curves of remanent magnetization under pressure versus pressure except for the goethite sample that retains 100% of its initial SIRM under 1.24 GPa (Fig. 4b), followed by a minor decrease in remanent magnetization (2%) after sample decompression. This is coherent with regard to its very high B cr value (>9 T after Rochette et al., 2005b). On the other hand, the maximum value of pressure demagnetization was obtained for a MD magnetite-bearing microdiorite which lost 84% of its initial SIRM under 1.24 GPa (sample est in Table 2 where ∆ is the percentage of demagnetization of the initial SIRM under 1.24 GPa).

Among the investigated pyrrhotite-bearing samples (Fig. 4a), PSD sample NWA 753 (Rumuruti chondrite) was found to be the most pressure resistant and synthetic samples of dispersed PSD pyrrhotite powder (pyr-a and pyr-b) were shown to be the most pressure sensitive. The monocrystal of predominately SD pyrrhotite (sample 98080) has a non-linear behavior resulting in a sharp decrease of IRM above 1.1 GPa. This could be related to a high anisotropy of the sample [START_REF] Louzada | Shock and Static Pressure Demagnetization of Pyrrhotite and Implications for the Martian Crust[END_REF] or/ and the approach to the magnetic phase transition pressure. SD greigite samples lost 10% and 26% of their initial remanent magnetization under pressure, respectively.

In the group of hematite-bearing samples the most pressure sensitive was the synthetic sample of MD hematite that lost =58% of its initial SIRM under pressure, much more than SD hematite sample (=38%), PSD rhyolite (=27%), and radiolarites or jasper ([30; 35]%). SD and MD hematite samples have the same provenance (see 2.2): fine grain fraction of SD hematite powder was obtained by crushing the corresponding MD hematite powder (natural ore).

In the group of iron-and nickel iron-bearing samples the synthetic MD iron sample was found to be the most pressure sensitive (=83%) while the tetrataenite-bearing samples (Saratov and Bensour-j meteorites) are very resistant to pressure (~3%). It is interesting to mention that within the group of Fe 1-y Ni y -bearing samples pressure demagnetization degree decreases almost A c c e p t e d M a n u s c r i p t 13 linearly with increasing Ni-content (y=0% for pure iron, y~10% for kamacite, y~30% for taenite and y~50% for tetrataenite). Additional work is needed to confirm this observation.

In the group of titanomagnetite-bearing samples (Fig. 4d) we investigated samples with different ulvospinel content (x) ranging from x=0.03 (ona12 and spi3301) to x=0.63 (kil-2).

Pressure sensitivity of titanomagnetites does not seem to be dependent on ulvospinel content alone. For instance, samples pd6-2-4 and au1b have very different x values (x=0.42 and x=0.09, respectively) but have almost the same pressure demagnetization degree under 1.24 GPa.

The group of magnetite-bearing samples shows the most diversified pressure demagnetization behavior and the highest dispersion in  under 1.24 GPa with values ranging from 5 to 84%. As seen from Fig. 4e-f, some pressure demagnetization curves are almost indistinguishable (e.g., samples bf8703, NWA 998 and iro8), in agreement with their similar grain size (PSD, see Table 1) and magnetic hardness (B cr  [41, 61] mT). Their different geological origin (bf8703 is a granite and iro8 is an ignimbrite and NWA 998 is a mafic cumulate) does not seem to play any significant role with respect to pressure sensitivity.

The values of median destructive field (MDF i in mT) and median destructive stress (MDS in GPa) of SIRM, needed to remove one half of the initial SIRM by alternating field and by pressure, respectively, are presented in Tables 12. Most of MDS values were obtained by linear extrapolation of the linear parts of pressure demagnetization curves. Such values are indicated in italic in Table 2. This computation assumes that pressure demagnetization curves above 1.24 GPa remain linear, and that the samples do not undergo any phase transitions at pressures lower than MDS. It is not the case for pyrrhotite that undergoes a high-pressure magnetic transition and loses 100% of its remanence at about 2.8 GPa [START_REF] Rochette | High pressure magnetic transition in pyrrhotite and impact demagnetization on Mars[END_REF]. On average, hematite-bearing samples have the lowest MDS values situated in 1.1 to 2.3 GPa range (see Table 2).

A c c e p t e d M a n u s c r i p t 14 As it is seen from Fig. 4, there is a variety of behaviors with respect to pressure demagnetization. In order to characterize different shapes of pressure demagnetization curves we introduced the following coefficient:

α=[(SIRM 0 -IRM p1 )/(SIRM 0 -IRM p2 )]/[p 1 /p 2 ] ( 1
)
where SIRM 0 is the initial SIRM of the sample in the pressure cell before pressure application, IRM p1 is IRM under p 1 =0.46 GPa and IRM p2 is IRM of the sample under p 2 =1.24 GPa. Pressure demagnetization curve has concave shape if  >1, convex shape if  <1 and linear shape if

1.
As seen from Table 2 and Fig. 4, pressure demagnetization curves of magnetite-and titanomagnetite-bearing sample have mostly linear or concave shapes (average α=1.4 for titanomagnetites and 1.5 for magnetites) whereas those of pyrrhotites (except synthetic samples pyr-a and pyr-b) have slightly convex shape (average α=0.8 without considering synthetic pyrrhotites) due to the existence of a 2.8 GPa phase transition [START_REF] Rochette | High pressure magnetic transition in pyrrhotite and impact demagnetization on Mars[END_REF]. Curves for nickel iron-bearing meteorites have both concave and convex shapes and those for hematitebearing samples have mostly concave or linear shapes (with average =1.4).

For iron a high pressure ferromagnetic to paramagnetic phase transition is known to take place at 13 GPa (e.g., [START_REF] Dickinson | Shock magnetism in fine particle iron[END_REF]: at this pressure point -iron (bodycentered cubic structure) is transformed into ε-iron (hexagonal close packed structure), which is known to be non-magnetic. In our study pressure demagnetization curve for the sample of synthetic iron has strongly concave shape; furthermore the sample has already lost up to 83% of its initial SIRM under 1.24 GPa. Therefore, an iron sample will most likely lose 100% of its magnetic remanence far below 13 GPa without undergoing any magnetic phase transition.

Discussion

In order to check the effect of repeated load (to the same pressure level) on the magnetic remanence, several cycles from 0 to 1.24 GPa were carried out on some of investigated sample A c c e p t e d M a n u s c r i p t 15 (see Fig. 8). Throughout repeated loads the remanent magnetization under pressure (Fig. 8a) or upon pressure release (Fig. 8b) always shows a slight decrease as a function of number of cycles, consistent with previous works (e.g., [START_REF] Pozzi | Magnetic properties of oceanic basaltseffects of pressure and consequences for the interpretation of anomalies[END_REF][START_REF] Gilder | Static stress demagnetization of single and multidomain magnetite with implications to meteorite impacts[END_REF][START_REF] Bezaeva | Pressure demagnetization of the Martian crust: Ground truth from SNC meteorites[END_REF]. This is not linked to a longer application of maximum pressure as it was shown above (see 3.1) that hydrostatic pressure demagnetization effect is time-independent.

Contrary to the repeated application of the same pressure demagnetization step, repeated application of the same alternating field demagnetization step does not produce a further demagnetization effect of magnetic remanence. This shows that pressure demagnetization and alternating field demagnetization have different physical nature and can not be directly compared in terms of process. Samples that have similar alternating field demagnetization behavior of SIRM can have very different pressure demagnetization behavior (e.g., samples spi3301 and bb, see Fig. 4d, 4j; or samples bf3201 and est, see Fig. 4e,4k). Therefore, it is not possible to link a certain pressure value to AF value giving the same degree of demagnetization. AF demagnetizes all grains with coercitivities under a given value whereas pressure demagnetizes grains within a large spectrum of coercivities. Moreover, the degree of demagnetization can vary over the coercivity spectrum of a sample (Gattacceca et al., 2007a).

To represent how pressure demagnetization affects different coercivity fractions of the IRM, we introduced in [START_REF] Bezaeva | Pressure demagnetization of the Martian crust: Ground truth from SNC meteorites[END_REF] the demagnetization loss :

(B)=[IRM B (B)-IRM BP (B)]/IRM B (B), (2)
as a function of alternating field B (in mT). IRM B and IRM BP correspond to IRM values before and after pressure application, respectively, so that IRM B (B) is a curve of AF demagnetization of initial SIRM and IRM BP (B) is a curve of AF demagnetization of residual IRM after decompression from 1.24 GPa (IRM PR ). Curves of (B) for Martian meteorites (NWA 1068, NWA 998, LA), basalt ba and rhyolite rb7a were presented in Bezaeva et al. (2007, Fig. 3). Table 2 gives (in %) the values of  1 =1-(B=30 mT)/(B=0 mT), characterizing the shape of (B) curves.

 1 =0 ( 1 =1
) if demagnetization affects only grains with coercivity below (above) 30 mT. In all A c c e p t e d M a n u s c r i p t 16 cases  decreases with increasing AF, indicating that pressure preferentially demagnetizes the lower coercivity fractions (consistent with previous works, e.g., [START_REF] Pearce | On pressure demagnetization[END_REF] but in some cases high coercivity fractions may also be affected. As seen from Table 2, the average  1 value for hematite-bearing samples is of (94)%, so mostly low coercivity fractions are affected by pressure demagnetization. For magnetite and titanomagnetite families high coercivity fractions (above 30 mT) are also affected: the average  1 is of (6927)% for magnetites and (6910)% for titanomagnetites. For pyrrhotite-and nickel iron-bearing samples average  1 is of (5423)% and (5329)%, respectively.

In most experiments pressure release resulted in further changes of the residual magnetic remanence comparing to the under-pressure level (1.24 GPa). Relative changes in magnetic remanence upon decompression (δ in %) are presented in Table 2. Different magnetic mineralogies have different response to decompression resulting in either further increase (positive δ) of or further decrease (negative δ) in residual remanence (or, more rarely, no changes). An increase of residual IRM upon decompression was observed for pyrrhotite-bearing samples (1<δ<8%) and hematite-bearing samples (0<δ<9%). No significant effect was observed on metal, except for the sample of synthetic iron that shows 6% decrease in IRM after pressure release. Ti-rich and Ti-poor(free) magnetites seem to have different behavior of magnetic remanence upon decompression, so they were separated into two groups. Changes in magnetic remanence for Ti-free and Ti-poor magnetites (with 0.03x0.09) are in the range -7 to +6%; mostly decrease or negligible increase in remanence are observed apart for synthetic magnetites and osb10b sample. Titanomagnetites with high ulvospinel content (x>0.4) have generally higher δ (in the 2 to 19% range). Remarkably high δ values are observed for ba and kil-2 samples (13% and 19% recovery of magnetic remanence upon pressure release, respectively). It is interesting to note that both samples are characterized by the highest values of ulvospinel content between all investigated titanomagnetite-bearing samples.

A c c e p t e d M a n u s c r i p t

17

The question of a key parameter that may be responsible for the pressure resistance of different materials has already been put forward in previous works [START_REF] Pearce | On pressure demagnetization[END_REF][START_REF] Kletetschka | Pressure effects on Martian crustal magnetization near large impact basins[END_REF]. [START_REF] Pearce | On pressure demagnetization[END_REF] proposed a simple correlation between coercivity and the effect of pressure demagnetization. [START_REF] Kletetschka | Pressure effects on Martian crustal magnetization near large impact basins[END_REF] suggested that the efficiency of pressure demagnetization is proportional to the logarithm magnetic coercivity (B c ), irrespective of magnetic mineralogy.

We have already shown that the pressure resistance of samples is tightly controlled by their magnetic mineralogy (see above). By analyzing our experimental data, we concluded that for a given magnetic mineralogy, the key proxy for sensitivity of samples to hydrostatic pressure demagnetization is B cr rather than B c . Indeed, for each magnetic mineralogy pressure resistance is roughly proportional to ln(B cr ) (Fig. 9). Synthetic samples of magnetite and iron were not considered for this approximation; synthetic samples of hematite were taken into consideration due to natural provenance of the used hematite powder (natural ore) (Fig. 9). Residual IRM under pressure p 1 =0.46 GPa or p 2 =1.24 GPa, normalized to initial SIRM (IRM p ), has been fitted using the following equation:

IRM px /SIRM = a x ln(B cr ) + b x , (3)
where x=1 for p 1 and x=2 for p 2 ; a x and b x coefficients as well as linear correlation coefficients r x between IRM px /SIRM and B cr (in mT) are given in Table 3. Strictly speaking B cr value should be dimensionless to equilibrate the equation (as the left part of the equation is dimensionless); this can be easily reached by normalizing the B cr value to 1 mT. Linear correlation coefficients r x * between IRM px /SIRM and B c (in mT) are also given in Table 3. Thus, based on the relative values of r x and r x * (see Table 3), B cr appears to be a better proxy than B c for all investigated magnetic mineralogies: hematite, pyrrhotite, nickel iron, Ti-magnetite and magnetite, and both pressure values (p 1 and p 2 ). This could be expected as magnetic remanence should be more directly related to remanent rather than induced properties. The highest scatter is observed for magnetite and titanomagnetite-bearing samples (Fig. 9d), pointing to a complex behaviour of magnetic remanence versus pressure for this magnetic mineralogy. In Fig. 10 efficiency of magnetites and Ti-magnetites is clearly grain-size dependent and increases with increasing grain size (decreasing M rs /M s ). Pressure demagnetization degree of MD magnetites in our study is consistent with the results of [START_REF] Gilder | Static stress demagnetization of single and multidomain magnetite with implications to meteorite impacts[END_REF], who showed that MD magnetite retained 66% of its initial SIRM under 1.40 GPa. We did not investigate Ti-free SD magnetite samples that would have allowed confirmation of the surprising result by [START_REF] Gilder | Static stress demagnetization of single and multidomain magnetite with implications to meteorite impacts[END_REF] that SD magnetite is more pressure sensitive than MD magnetite despite its higher B cr .

Metal and Ti-free magnetite have the highest intrinsic pressure resistivity (after goethite), but become much more pressure sensitive in their multidomain state. A synthetic graph for all magnetic mineralogies under p 2 =1.24 GPa is presented in Fig. 11. The B cr value corresponding to the threshold, above which no further pressure demagnetization could be produced under 1.24

GPa (called B th ), was determined for each magnetic mineralogy (Table 3). The highest value of values, magnetite and Ti-magnetite appear as pressure sensitive as nickel iron but more pressure resistant than pyrrhotite (in the coercivity window B cr >33 mT). Goethite was found to be the most pressure resistant mineral. Based on the two greigite samples, this mineral appears to be more sensitive to pressure than magnetite.

Conclusions

This study gives an overview of the sensitivity to pressure demagnetization (by a purely hydrostatic load up to 1.24 GPa) of geological and extraterrestrial materials as well as synthetic samples with a variety of magnetic mineralogies: magnetite and titanomagnetite, pyrrhotite, greigite, hematite, goethite as well as iron and iron-nickel alloys. Magnetic remanence under pressure and upon decompression was investigated using a non-magnetic high pressure cell of piston-cylinder type together with a high sensitivity remanence magnetometer (SQUID magnetometer).

After showing the good reproducibility of our experiments, we demonstrated the additivity of partial pressure demagnetization. It was found that under hydrostatic pressure of 1.24 GPa, applied in a low magnetic field (<5 μT), rock and synthetic samples lost irreversibly up to 84% of their initial saturation isothermal remanent magnetization without any changes in their intrinsic magnetic properties. Decompression resulted in further changes in IRM (decrease or increase up to 19% with respect to the initial SIRM value before compression) but decompressed residual IRM never reached SIRM value before compression.

It was shown that hydrostatic pressure demagnetization is not time-dependent, but repeated application of the same pressure level leads to further demagnetization. The main factor controlling the stability of remanent magnetization versus hydrostatic pressure is magnetic mineralogy. For a given magnetic mineralogy the resistance to pressure is roughly proportional Thus, based on our experimental investigations, we propose a phenomenological model allowing to quantify the pressure demagnetization effect (Fig. 11, Table 3): after identification of the main magnetic mineralogy of the sample and by measuring its B cr value, it is possible to roughly estimate the resistance of its magnetic remanence to a purely hydrostatic load of 0.46 and 1.24 GPa. For magnetite and titanomagnetite such a model must be refined by considering the magnetic domain state and the substitution rate x (Fig. 10). The pressure demagnetization effect was shown to be grain-size sensitive: the efficiency of pressure demagnetization typically increases with increasing grain size. m is mass (in mg). T C is Curie temperature (in C); theoretical value when italized.  0 is low field magnetic susceptibility (in 10 -6 m 3 /kg). MDF i is median destructive field of IRM (in mT): 9T IRM for hematite-bearing and goethite samples and 3T IRM for all other samples. B c and B cr are magnetic coercivity and coercivity of remanence, respectively (in mT). M s and M rs are saturation and saturation remanent magnetization (in 10 -3 Am 2 /kg). "Ti-magn" is titanomagnetite. For synthetic samples numbers between brackets (1 st column) correspond to estimated mass and volume concentrations of powders in the sample respectively (in %) and mass values correspond to the mass of the whole samples of dispersed powder with epoxy resin matrix. In the second column the information between brackets A c c e p t e d M a n u s c r i p t All SIRM and IRM values are expressed in Am 2 ; SIRM 1 is the initial SIRM of the sample (before compression) in the pressure cell (at zero pressure); IRM p2 is residual IRM under maximum pressure p 2 =1.24 GPa (inside the cell);  (in %) is pressure demagnetization degree under p 2 ; IRM PR is residual IRM after decompression from p 2 and extraction of the sample from the cell. δ (in %) corresponds to changes in IRM upon decompression from p 2 (IRM decreases when δ is negative and increases when δ is positive).

SIRM 2 is SIRM acquired after decompression from p 2 and new saturation; MDS is median destructive stress (in GPa); it is italized when linear extrapolation of pressure demagnetization curves was used for MDS determination. In the right column numbers between brackets correspond to  0 values at 0 mT (in %).

α=[(SIRM 0 -IRM p1 )/(SIRM 0 -IRM p2 )]/[p 1 /p 2 ], where p 1 =0.46 GPa;  1 =1-(30 mT)/ 0 , where (B)=[IRM B (B)-IRM BP (B)]/IRM B (B),  0 =(0 mT) and B is alternating field (see text).

( 1) Values of remanent magnetization correspond to the NRM of the cell (that was never saturated);

(2) The initially compressed magnetization was TRM; (3) The initially compressed magnetization was NRM A c c e p t e d M a n u s c r i p t corresponds to reference(s) where samples are described and/ or their magnetic properties measured. (0) this study;

(1-6) Gattacceca et al., 2003, 2004, 2006, 2007a, 2007b and 2008, respectively; (7-8) Louzada et al., 2007, 2008; (9-15) A c c e p t e d M a n u s c r i p t All SIRM and IRM values are expressed in Am 2 ; SIRM 1 is the initial SIRM of the sample (before compression) in the pressure cell (at zero pressure); IRM p2 is residual IRM under maximum pressure p 2 =1.24 GPa (inside the cell);  (in %) is pressure demagnetization degree under p 2 ; IRM PR is residual IRM after decompression from p 2 and extraction of the sample from the cell. δ (in %) corresponds to changes in IRM upon decompression from p 2 (IRM decreases when δ is negative and increases when δ is positive).

SIRM 2 is SIRM acquired after decompression from p 2 and new saturation; MDS is median destructive stress (in GPa); it is italized when linear extrapolation of pressure demagnetization curves was used for MDS determination. In the right column numbers between brackets correspond to  0 values at 0 mT (in %).

α=[(SIRM 0 -IRM p1 )/(SIRM 0 -IRM p2 )]/[p 1 /p 2 ], where p 1 =0.46 GPa;  1 =1-(30 mT)/ 0 , where (B)=[IRM B (B)-IRM BP (B)]/IRM B (B),  0 =(0 mT) and B is alternating field (see text).

( 1) Values of remanent magnetization correspond to the NRM of the cell (that was never saturated);

(2) The initially compressed magnetization was TRM; (3) The initially compressed magnetization was NRM Pearson's correlation is determined as a ration of covariance of the two variables to the product of their standard deviation. B c is magnetic coercivity. R 2 x is the coefficient of determination (R-squared value), it ranges from 0 to 1 and reveals how closely the logarithmic fit IRM px =f(B cr ) corresponds to the actual data. The calculated R 2 is not an adjusted R-squared value (the transformed regression model is used for calculations). B th is threshold B cr above which no demagnetization under p 2 is possible. B th = e (1-b)/a .

Figure

  Figure 4a to 4f displays curves of IRM under pressure, normalized to initial SIRM, versus

  or R-squared values (R 2 x ) reflecting the confidence of a chosen approximation were calculated for logarithmic (3) and linear approximations IRM px /SIRM=f(B cr ). It was shown that R 2x was always higher for logarithmic than for linear approximation (except for pyrrhotite under p 1 , where both logarithmic and linear approximations have very low R-squared values).

  IRM p2 /SIRM is plotted versus M rs /M s for all natural magnetite-and titanomagnetite-bearing samples. M rs and M s are saturation remanent and induced magnetizations, respectively. As shown in Fig. 10, pressure demagnetization of (Ti-)magnetites is controlled by both ulvospinel content (vertical shift of points on the graph) and magnetic domain structure (horizontal shift of points on the graph). Three sample groups can be delineated: samples with x[0; 0.1] (a); samples with x[0.40; 0.55] (b) and samples with x0.55 (c). The last group is represented by only two samples (LA and kil-2). Samples from the three groups lie along three separate trends. As it is seen from Fig. 10, pressure demagnetization

B

  th corresponds to hematite, which appears to be the most pressure sensitive material in agreement with the lowest MDS values found for hematite-bearing samples. Based on other B th

  B cr ), where B cr is the coercivity of remanence (reflecting the magnetic hardness of the sample).

Figure 1 .

 1 Figure captions:

Figure 2 .

 2 Figure 2. Example of correction of the pressure demagnetization curve of hematite-bearing

Figure 3 .

 3 Figure 3. Residual isothermal remanent magnetization under pressure (normalized to initial

Figure 4

 4 Figure 4 (a-f): Residual isothermal remanent magnetization under pressure (normalized to initial

Figure 5 .

 5 Figure 5. Residual isothermal remanent magnetization under pressure (normalized to initial

Figure 6 .

 6 Figure 6. Residual remanent magnetization under pressure (normalized to its initial value of

Figure 7 .

 7 Figure 7. Alternating field (AF) demagnetization curves of initial SIRM, residual IRM after

Figure 8 .

 8 Figure 8. (a) Residual remanent magnetization under 1.24 GPa versus number of compression

Figure 9 .

 9 Figure 9. Residual remanent magnetization (normalized to initial SIRM) under p 1 =0.46 GPa

Figure 10 .

 10 Figure 10. Residual remanent magnetization (normalized to initial SIRM) under 1.24 GPa

Figure. 11 .

 11 Figure. 11. Synthetic graph of residual remanent magnetizations (normalized to initial SIRM)

Figure

  Figure(s)

Table 1

 1 Main magnetic properties of investigated samples.

	TABLES
	Sample	Main magnetic carrier and ref.	m	T C	 0	MDF i	B cr	B cr /B c	M s	M rs /M s
	METEORITES
	Martian meteorites
	Los Angeles	Ti-magn (11, 13)	94.3	150	10.3	31	55	1.41	1232	0.49
	NWA 1068	pyrrhotite (13, 11)	240.2	325	1.3	100	134	1.54	212	0.47
	NWA 998	magnetite (13, 11)	20.4	575	4.3	25	41	1.73	713	0.24
	Ordinary chondrites
	Bensour-6a	taenite (0, 1)	162.0	450	154.0	18	33	35.8	3187	0.003
	Bensour-i	tetrataenite (0, 1)	23.7	580	-	24	64	1.62	1963	0.16
	Bensour-j	tetrataenite (0, 1)	102.4	580	46.8	25	64	1.62	1963	0.16
	Pultusk	kamacite (0)	64.2	760	600.0	10	27	26.0	303	0.007
	Saratov	tetrataenite (0)	54.1	580	633.0	70	162	35.8	16540	0.03
	Rumuruti chondrites
	NWA 753	pyrrhotite (0, 15)	211.9	325	8.7	108	134	1.68	93	0.41
	TERRESTRIAL ROCKS AND MINERALS
	Basalts
	ba	Ti-magn (9)	59.5	50	16.1	15	18	3.54	622	0.16
	ba (TRM)	Ti-magn (9)	405.0	50	16.1	15	18	3.54	622	0.16
	bb-a	Ti-magn (0, 6)	327.8 60/150	16.1	8	14	2.29	866	0.18
	bb-a (TRM)	Ti-magn (0, 6)	327.3 60/150	16.1	8	14	2.29	866	0.18
	bb-b (NRM)	Ti-magn (0, 6)	392.3 60/150	16.1	9	14	2.29	866	0.18
	be-3a	Ti-magn (3, 10)	267.7	245	10.8	13	18	2.57	1700	0.13
	pd6-2-1	Ti-magn (0, 8)	30.8	280	109.0	15	27	1.86	1827	0.23
	pd6-2-4	Ti-magn (0, 8)	51.2	280	131.0	11	21	1.79	1833	0.21
	Andesites
	au1b	Ti-magn (0, 5)	226.2	530	121.4	8	16	4.24	1828	0.03
	cug1b	Ti-magn (0, 5)	175.5	550	88.7	7	14	4.65	1048	0.03
	osb10b	magnetite (0, 5)	125.3	575	9.74	40	69	2.10	139	0.34
	Ignimbrites
	spi3301	Ti-magn (0, 5)	302.8	560	73.0	9	16	2.81	545	0.11
	ona12	Ti-magn (0, 5)	109.4	560	45.4	24	44	2.86	787	0.12
	spi3903	magnetite (0, 5)	233.0	575	54.1	16	27	2.79	621	0.13
	mtd5b	magnetite (0, 5)	162.3	575	93.7	22	40	2.37	1356	0.18
	iro8	magnetite (0, 5)	161.0	575	13.6	31	48	2.03	151	0.27
	Rhyolites
	rb7a	hematite (16)	238.0	650	0.03	>150	406	3.56	5	0.48
	Radiolarites
	radiol-1	hematite (0)	333.5	675	1.4	>150	390	1.21	67	0.68

Table 2 .

 2 Synthesis of results on pressure demagnetization experiments.

	Sample	SIRM 1	IRM p2		IRM PR	δ	SIRM 2 MDS	α	 1
	pressure cell (1)	0.0343	0.0433	-	-	-	-	-	-	-
	METEORITES
	Los Angeles	11.9	9.11	23	3.08	3.1	9.20	2.5	1.3	21(67)
	NWA 1068	25.3	21.1	17	19.8	3.2	24.6	3.7	0.8	36(20)
	NWA 998	7.97	7.47	6	6.25	-0.4	7.36	9.5	1.1	69(15)
	Bensour-6a	1.36	1.10	19	1.01	-3.7	1.57	3.1	0.8	80(36)
	Bensour-i	6.81	6.60	3	6.49	2.1	6.28	22.4	1.3	-
	Bensour-j	32.0	30.7	4	30.6	-0.9	32.0	13.3	1.0	46(4)
	Pultusk	15.3	10.2	33	9.89	0	16.0	1.8	1.3	98(38)
	Saratov	20.3	19.6	3	18.5	0.5	-	18.9	0.8	20(16)
	NWA 753	7.80	7.05	10	6.94	1.3	7.65	6.5	1.0	44(9)
	TERRESTRIAL ROCKS AND MINERALS
	ba	6.78	3.12	54	3.56	12.5	6.13	1.07	1.7	29(42)
	ba (2)	5.52	3.64	34	-	10.9	-	1.7	1.0	-
	bb-a	48.3	24.9	48	25.7	1.9	52.5	1.26	1.4	95(51)
	bb-a (2)	2.12	1.11	48	1.63	5.1	52.5	1.26	1.5	4(97)
	bb-b (3)	1.62	0.96	40	0.91	9.1	2.12	1.5	1.0	72(57)
	be-3a	74.9	32.1	57	36.9	6.8	71.7	1.08	-	91(48)
	pd6-2-1	12.1	8.51	30	8.91	3.9	12.1	2.6	1.5	99(26)
	pd6-2-4	18.5	11.1	40	10.6	2.2	19.1	1.6	1.4	97(45)
	au1b	13.1	7.55	42	7.62	0.3	12.7	1.4	1.3	-
	cug1b	5.62	3.04	46	2.62	-6.9	5.92	1.29	1.2	88(56)
	osb10b	5.64	3.89	31	4.22	5.7	5.91	2.6	1.6	59(29)
	spi 3301	19.1	13.2	31	10.4	-7.3	16.5	1.9	1.3	81(37)
	ona12	10.4	8.91	14	8.72	0.2	10.4	4.2	1.1	89(16)
	spi3903	18.1	12.4	31	12.1	1.7	18.3	2.7	1.6	75(34)
	mtd5b	38.0	33.0	13	37.5	-3.9	37.2	5.1	1.4	75(19)
	iro8	6.37	6.05	5	5.73	-0.2	6.20	12.4	0.8	88(8)
	rb7a	0.55	0.40	27	0.44	8.9	0.60	2.1	1.2	14(27)
	radiol-1	17.9	11.9	34	11.7	0.0	18.7	2.3	1.7	3(38)
	radiol-2	13.2	8.53	35	9.45	8.1	13.5	2.0	1.7	6(30)
	bf0703	9.47	8.22	13	7.80	-2.5	8.47	4.7	1.1	-
	bf3201	2.76	0.95	66	0.75	-4.6	2.72	0.54	1.8	65(73)
	bf9804	1.41	0.77	46	0.73	-3.9	1.54	1.5	1.7	66(53)
	bf8703	4.14	3.92	5	3.79	-1.7	4.13	11.9	1.2	61(8)
	est	8.09	1.32	84	0.98	-0.5	8.17	0.19	2.4	43(88)
	sb3a	15.5	12.2	22	11.8	5.8	20.3	3.7	0.9	25(42)

Table 3 .

 3 Parameters related to pressure sensitivity of different magnetic mineralogies.Residual isothermal remanent magnetization under pressure normalized to its initial SIRM (IRM p ) can be calculated for p 1 =0.46 GPa and p 2 =1.24 GPa using the following empirical equation: IRM px /SIRM = a x ln(B cr ) + b x , where x=1 for p 1 and x=2 for p 2 ; B cr is coercivity of remanence in mT. r x and r x * are Pearson correlation coefficients indicating the strength of linear dependences between IRM px /SIRM and B cr as well as IRM px /SIRM and B c , respectively.Pearson's correlation is determined as a ration of covariance of the two variables to the product of their standard deviation. B c is magnetic coercivity. R 2 x is the coefficient of determination (R-squared value), it ranges from 0 to 1 and reveals how closely the logarithmic fit IRM px =f(B cr ) corresponds to the actual data. The calculated R 2 and R 2 x are correlation coefficient between IRM p and B cr and determination coefficient of logarithmic fits IRM px =f(B cr ). r x * is correlation coefficient for the case when B cr is replaced by B c , where B c is magnetic coercivity. is not an adjusted Rsquared value (the transformed regression model is used for calculations). B th is threshold B cr above which no demagnetization under p 2 is possible. B th = e (1-b)/a .

				p 1 =0.46 GPa							p 2 =1.24 GPa
	Magnetic
		a 1	b 1	r 1	r 1	*	R 2	1	a 2	b 2	r 2	r 2	*	R 2	2	B th (mT)
	mineralogy
	pyrrhotite	0.005	0.90	0.38	0.39	0.01	0.06	0.56	0.73	0.71	0.63	110036	Formatted: English (U.S.)
	nickel iron	0.14	0.42	0.88	0.81	0.82	0.35	-0.49	1.00	0.99	1.00	701	Formatted: English (U.S.)
	hematite	0.04	0.60	0.53	0.08	0.43	0.11	0.04	0.90	0.59	0.93	6100073
	(Ti)magnetite 0.15	0.34	0.65	0.57	0.47	0.27	-0.22	0.74	0.66	0.63	10099
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Table 1

 1 Main magnetic properties of investigated samples.

	Sample	Main magnetic carrier and ref.	m	T C	 0	MDF i	B cr	B cr /B c	M s	M rs /M s
	METEORITES
	Martian meteorites
	Los Angeles	Ti-magn (11, 13)	94.3	150	10.3	31	55	1.41	1232	0.49
	NWA 1068	pyrrhotite (13, 11)	240.2	325	1.3	100	134	1.54	212	0.47
	NWA 998	magnetite (13, 11)	20.4	575	4.3	25	41	1.73	713	0.24
	Ordinary chondrites
	Bensour-6a	taenite (0, 1)	162.0	450	154.0	18	33	35.8	3187	0.003
	Bensour-i	tetrataenite (0, 1)	23.7	580	-	24	64	1.62	1963	0.16
	Bensour-j	tetrataenite (0, 1)	102.4	580	46.8	25	64	1.62	1963	0.16
	Pultusk	kamacite (0)	64.2	760	600.0	10	27	26.0	303	0.007
	Saratov	tetrataenite (0)	54.1	580	633.0	70	162	35.8	16540	0.03
	Rumuruti chondrites
	NWA 753	pyrrhotite (0, 15)	211.9	325	8.7	108	134	1.68	93	0.41
	TERRESTRIAL ROCKS AND MINERALS
	Basalts
	ba	Ti-magn (9)	59.5	50	16.1	15	18	3.54	622	0.16
	ba (TRM)	Ti-magn (9)	405.0	50	16.1	15	18	3.54	622	0.16
	bb-a	Ti-magn (0, 6)	327.8 60/150	16.1	8	14	2.29	866	0.18
	bb-a (TRM)	Ti-magn (0, 6)	327.3 60/150	16.1	8	14	2.29	866	0.18
	bb-b (NRM)	Ti-magn (0, 6)	392.3 60/150	16.1	9	14	2.29	866	0.18
	be-3a	Ti-magn (3, 10)	267.7	245	10.8	13	18	2.57	1700	0.13
	pd6-2-1	Ti-magn (0, 8)	30.8	280	109.0	15	27	1.86	1827	0.23
	pd6-2-4	Ti-magn (0, 8)	51.2	280	131.0	11	21	1.79	1833	0.21
	Andesites
	au1b	Ti-magn (0, 5)	226.2	530	121.4	8	16	4.24	1828	0.03
	cug1b	Ti-magn (0, 5)	175.5	550	88.7	7	14	4.65	1048	0.03
	osb10b	magnetite (0, 5)	125.3	575	9.74	40	69	2.10	139	0.34
	Ignimbrites
	spi3301	Ti-magn (0, 5)	302.8	560	73.0	9	16	2.81	545	0.11
	ona12	Ti-magn (0, 5)	109.4	560	45.4	24	44	2.86	787	0.12
	spi3903	magnetite (0, 5)	233.0	575	54.1	16	27	2.79	621	0.13
	mtd5b	magnetite (0, 5)	162.3	575	93.7	22	40	2.37	1356	0.18
	iro8	magnetite (0, 5)	161.0	575	13.6	31	48	2.03	151	0.27
	Rhyolites
	rb7a	hematite (16)	238.0	650	0.03	>150	406	3.56	5	0.48
	Radiolarites
	radiol-1	hematite (0)	333.5	675	1.4	>150	390	1.21	67	0.68

Table (

 (

	s)
	Sample	Main magnetic carrier and ref.	m	T C	 0
				M a n u s c r i p t
	A c c e p t e d

MDF i B cr B cr /B c M s M rs /M s

  T C is Curie temperature (in C); theoretical value when italized.  0 is low field magnetic susceptibility (in 10 -6 m 3 /kg). MDF i is median destructive field of IRM (in mT): 9T IRM for hematite-bearing and goethite samples and 3T IRM for all other samples. B c and B cr are magnetic coercivity and coercivity of remanence, respectively (in mT). M s and M rs are saturation and saturation remanent magnetization (in 10 -3 Am 2 /kg). "Ti-magn" is

	radiol-2	hematite (0)	232.1	675	1.6	>150	347	1.30	74	0.69
	Granites
	bf0703	magnetite (0, 2)	304.6	575	11.6	39	68	2.98	165	0.15
	bf3201	magnetite (0, 2)	425.5	575	41.3	7	13	7.54	448	0.01
	bf9804	magnetite (0, 2)	453.3	575	8.8	13	25	6.01	112	0.03
	bf8703	magnetite (0, 2)	228.0	575	3.8	34	61	2.08	76	0.24
	Microdiorite
	est	magnetite (4)	320.0	575	13.5	8	19	10.3	128	0.01
	Schist
	sb3a	pyrrhotite (4)	243.3 280/320	0.3	17	31	1.29	91	0.50
	Jasper
	jas	hematite (0)	210.0	675	4.1	>150	289	1.38	154	0.70
	Obsidian
	kil-2	Ti-magn (0)	219.4	142	4.4	21	27	1.34	822	0.45
	Minerals (monocrystals)
	98080	pyrrhotite (0, 7)	429	325	11.8	4	4	0.88	1444	0.46
	127037	pyrrhotite (0, 7)	24.5	325	6.2	3	7	1.51	5696	0.15
	goethite	goethite (0, 14)	320.1	~120	0.3	>150 >9000	-	-	-
	Sediments
	greig-tw1	greigite (17)	5.0	330	970	59	80	1.38	19470	0.52
	greig-it1	greigite (18)	4.1	330	1408	52	73	1.45	29073	0.45
	SYNTHETIC SAMPLES of dispersed powders in epoxy resin
	pyr-a (20.5; 5.2)	pyrrhotite (0, 12)	33.5	325	175	7	17	1.73	-	0.21
	pyr-b (13.1; 3.4)	pyrrhotite (0, 12)	32.8	325	69.2	11	23	1.51	-	0.32
	magn1a (0.3; 0.1)	magnetite (0)	170.0	575	15.1	15	32	3.22	-	0.08
	magn6a (17.3; 4)	magnetite (0)	73.6	575	-	14	28	3.02	-	0.09
	magn6c (0.8; 0.2)	magnetite (0)	89.4	575	-	15	29	3.03	-	0.09
	magn6d (0.5; 0.1)	magnetite (0)	79.4	575	-	15	30	3.10	-	0.09
	iron-2 (2.5; 0.4)	iron (0)	229.5	753	281	8	22	13.8	-	0.004
	MD hem (2.1; 0.5)	hematite (0)	116.1	675	0.95	67	33	1.11	-	0.57
	SD hem	hematite (0)	155.6	675	-	>100	177	1.43	-	0.73
	m is mass (in mg).

titanomagnetite. For synthetic samples numbers between brackets (1 st column) correspond to estimated mass and volume concentrations of powders in the sample respectively (in %) and mass values correspond to the mass of the whole samples of dispersed powder with epoxy resin matrix. In the second column the information between brackets

Table 2 .

 2 Synthesis of results on pressure demagnetization experiments.

	Sample	SIRM 1	IRM p2		IRM PR	δ	SIRM 2 MDS	α	 1
	pressure cell (1)	0.0343	0.0433	-	-	-	-	-	-	-
	METEORITES
	Los Angeles	11.9	9.11	23	3.08	3.1	9.20	2.5	1.3	21(67)
	NWA 1068	25.3	21.1	17	19.8	3.2	24.6	3.7	0.8	36(20)
	NWA 998	7.97	7.47	6	6.25	-0.4	7.36	9.5	1.1	69(15)
	Bensour-6a	1.36	1.10	19	1.01	-3.7	1.57	3.1	0.8	80(36)
	Bensour-i	6.81	6.60	3	6.49	2.1	6.28	22.4	1.3	-
	Bensour-j	32.0	30.7	4	30.6	-0.9	32.0	13.3	1.0	46(4)
	Pultusk	15.3	10.2	33	9.89	0	16.0	1.8	1.3	98(38)
	Saratov	20.3	19.6	3	18.5	0.5	-	18.9	0.8	20(16)
	NWA 753	7.80	7.05	10	6.94	1.3	7.65	6.5	1.0	44(9)
	TERRESTRIAL ROCKS AND MINERALS
	ba	6.78	3.12	54	3.56	12.5	6.13	1.07	1.7	29(42)
	ba (2)	5.52	3.64	34	-	10.9	-	1.7	1.0	-
	bb-a	48.3	24.9	48	25.7	1.9	52.5	1.26	1.4	95(51)
	bb-a (2)	2.12	1.11	48	1.63	5.1	52.5	1.26	1.5	4(97)
	bb-b (3)	1.62	0.96	40	0.91	9.1	2.12	1.5	1.0	72(57)
	be-3a	74.9	32.1	57	36.9	6.8	71.7	1.08	-	91(48)
	pd6-2-1	12.1	8.51	30	8.91	3.9	12.1	2.6	1.5	99(26)
	pd6-2-4	18.5	11.1	40	10.6	2.2	19.1	1.6	1.4	97(45)
	au1b	13.1	7.55	42	7.62	0.3	12.7	1.4	1.3	-
	cug1b	5.62	3.04	46	2.62	-6.9	5.92	1.29	1.2	88(56)
	osb10b	5.64	3.89	31	4.22	5.7	5.91	2.6	1.6	59(29)
	spi 3301	19.1	13.2	31	10.4	-7.3	16.5	1.9	1.3	81(37)
	ona12	10.4	8.91	14	8.72	0.2	10.4	4.2	1.1	89(16)
	spi3903	18.1	12.4	31	12.1	1.7	18.3	2.7	1.6	75(34)
	mtd5b	38.0	33.0	13	37.5	-3.9	37.2	5.1	1.4	75(19)
	iro8	6.37	6.05	5	5.73	-0.2	6.20	12.4	0.8	88(8)
	rb7a	0.55	0.40	27	0.44	8.9	0.60	2.1	1.2	14(27)
	radiol-1	17.9	11.9	34	11.7	0.0	18.7	2.3	1.7	3(38)
	radiol-2	13.2	8.53	35	9.45	8.1	13.5	2.0	1.7	6(30)
	bf0703	9.47	8.22	13	7.80	-2.5	8.47	4.7	1.1	-
	bf3201	2.76	0.95	66	0.75	-4.6	2.72	0.54	1.8	65(73)
	bf9804	1.41	0.77	46	0.73	-3.9	1.54	1.5	1.7	66(53)
	bf8703	4.14	3.92	5	3.79	-1.7	4.13	11.9	1.2	61(8)
	est	8.09	1.32	84	0.98	-0.5	8.17	0.19	2.4	43(88)
	sb3a	15.5	12.2	22	11.8	5.8	20.3	3.7	0.9	25(42)

Table 3 .

 3 Parameters related to pressure sensitivity of different magnetic mineralogies.

p 1 =0.46 GPa p 2 =1.24 GPa

Magnetic mineralogy a 1 b 1 r 1 r 1 * R 2 1 a 2 b 2 r 2 r 2 * R 2 2 B th (mT)

  Residual isothermal remanent magnetization under pressure normalized to its initial SIRM (IRM p ) can be calculated for p 1 =0.46 GPa and p 2 =1.24 GPa using the following empirical equation: IRM px /SIRM = a x ln(B cr ) + b x , where x=1 for p 1 and x=2 for p 2 ; B cr is coercivity of remanence in mT. r x and r x * are Pearson correlation coefficients indicating the strength of linear dependences between IRM px /SIRM and B cr as well as IRM px /SIRM and B c , respectively.

	pyrrhotite	0.005	0.90	0.38	0.39	0.01	0.06	0.56	0.73	0.71	0.63	1100
	nickel iron	0.14	0.42	0.88	0.81	0.82	0.35	-0.49	1.00	0.99	1.00	70
	hematite	0.04	0.60	0.53	0.08	0.43	0.11	0.04	0.90	0.59	0.93	6100
	(Ti)magnetite 0.15	0.34	0.65	0.57	0.47	0.27	-0.22	0.74	0.66	0.63	100

Acknowledgements

This work was supported by the French Agence Nationale de la Recherche (project 05-JCJC-0133) and was partially funded by the CNRS-RFFI PICS program (grant №07-05-92165) while the stay of N.S. Bezaeva at CEREGE was funded by a research grant of the French Government (№2005814). We acknowledge K.L. Louzada (Harvard University, Cambridge, USA) for providing basalt samples from the Lonar crater (PD6-2-1, PD6-2-4). C. Francis (Harvard Museum of Natural History, Cambridge, USA) is acknowledged for providing the pyrrhottite samples 98080 and 127037. We are grateful to B. Zanda (Museum National d'Histoire Naturelle, Paris, France) and M.A. Nazarov (Vernadsky Institute for Geochemistry and Analytical Chemistry RAS, Moscow, Russia) for providing some meteorite samples. We thank A.P. Roberts (University of Southampton, Southampton, UK) and A.R. Muxworthy (Imperial College of London, London, UK) for providing greigite samples.

A c c e p t e d M a n u s c r i p t corresponds to reference(s) where samples are described and/ or their magnetic properties measured. (0) this study; (1-6) Gattacceca et al., 2003, 2004, 2006, 2007a, 2007b and 2008, respectively; (7-8) Louzada et al., 2007, 2008; (9-15) Rochette et al., 1993, 1998, 2001, 2003, 2005a, 2005b, 2008, respectively; (16) Vlag. et al., 1997; (17, 18) [START_REF] Horng | Inconsistent magnetic polarities between greigite-and pyrrhotite/ magnetite-bearing marine sediments from the Tsailiao-chi section, southwestern Taiwan[END_REF][START_REF] Jiang | Contradictory magnetic polarities in sediments and variable timing of neoformation of authigenic greigite[END_REF][START_REF] Florindo | Palaeomagnetism and rock magnetism in the upper Pliocene Valle Ricca (Rome, Italy) section[END_REF]Van Dongen et al., 2007}, respectively.