# Seismic structure of Martian mantle inferred from in situ sound velocity measurements

Paul Chauvigné, Geeth Manthilake, Denis Andrault, Julien Chantel, Emmanuel Gardés, Sylvie Demouchy, Fabrice Barou, Louis Hennet, Antoine Mathieu, L Henry, et al.

► To cite this version:

Paul Chauvigné, Geeth Manthilake, Denis Andrault, Julien Chantel, Emmanuel Gardés, et al. Seismic structure of Martian mantle inferred from in situ sound velocity measurements. Physics of the Earth and Planetary Interiors, 2025, 365, pp.107378. 10.1016/j.pepi.2025.107378. hal-05076615

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# Seismic structure of Martian mantle inferred from *in situ* sound velocity measurements

P. Chauvigné¹, ⁵*, G. Manthilake¹, D. Andrault¹, J. Chantel², E. Gardés¹, S. Demouchy¹, ³, F. Barou³, L. Hennet⁴, A. Mathieu¹, L. Henry⁵, N. Guignot⁵, N. Bolfan-Casanova¹

¹ Laboratoire Magmas et Volcans. Université Clermont Auvergne. CNRS. IRD. OPGC;
63000 Clermont-Ferrand, France.

² Univ. Lille. CNRS. INRAE. Centrale Lille. UMR 8207 - UMET - Unité Matériaux et Transformations ; 59000 Lille. France.

³ Geosciences Montpellier. UMR5243. CNRS and Université de Montpellier; Montpellier.
France

⁴ Interfaces, Confinement, Matériaux et Nanostructures. CNRS-Université d'Orléans. UMR
7374 ; Orléans Cedex 2. France

⁵ Synchrotron SOLEIL ; 91192 Gif-sur-Yvette. France.

* Now at ⁵

*Corresponding author: Paul Chauvigné (paul.chauvigne@synchrotron-soleil.fr)

## Highlights:

* *In situ* sound velocity measurements are performed at high pressure

* Two pyrolytic compositions different in iron content are used (Fe#=0.17 and 0.27)

* Seismic profiles along several possible Martian areotherms are modeled

* The iron content of the Martian mantle should be lower than previously estimated

* The discontinuity at ~1000 km of depth is induced by the olivine to wadsleyite transformation

## Keywords:

Mars – Interior – Experimental petrology – Seismic velocities – Mantle structure

## Abstract

We report elastic wave velocity (Vp: P-Waves velocity, and Vs: S-wave velocities) measurements of iron-enriched pyrolite aggregates by a combination of ultrasonic interferometry, synchrotron radiation techniques and multi-anvil apparatus experiments. We carried out eight synchrotrons *in situ* experiments on two different bulk pyrolite compositions with Fe/(Fe+Mg) ratio (Fe#) of 0.17 and 0.27, over a range of pressures from 3 to 16 GPa and temperatures up to 1000 °C. These data are complemented by seven laboratory *ex situ* experiments between 5 and 16 GPa and temperatures varying from 900 °C to 1300 °C, using the same starting compositions. Based on these results and previous data on pyrolite composition, we modeled the effect of iron on seismic profiles along several possible Martian areotherms.

If, as reported in previous works, the Martian mantle had a very high Fe#, for instance Fe# = 0.27, it would be composed of four layers separated by three seismic discontinuities. They would be located at depths of ~620 km (ΔVp of +8.25%; ΔVs of +6.72%), ~880 km (ΔVp of +2.35%; ΔVs of 0.03%) and ~1130 km (ΔVp of -17.88%; ΔVs of -13.05%).
However, since there is no major discontinuity inferred by InSight's seismic data at depths shallower than 1000 km, we conclude that the Martian mantle presents a Fe# significantly lower than 0.20. For the Fe#0.17 composition, our experiments show a single discontinuity (ΔVp of +5.62%; ΔVs of +10.44%) induced by the transition from olivine to wadsleyite at ~ 980 km depth, which is compatible with the seismic data.

**Plain Language Summary**

The geochemical and geophysical data available for Mars suggest an interior enriched in iron, but this information alone is insufficient to constrain the mineralogy and the internal structure of the Martian mantle. In this study, we determine experimentally the elastic wave velocities (Vp: P-Waves velocity, and Vs: S-wave velocities) of two different iron-rich mantle compositions, with Fe# ratios (=Fe/(Mg+Fe)) of 0.17 and 0.27, in a range of pressure and temperature covering the Martian mantle conditions. We then calculate the seismic wave velocity profiles expected in the Mars interior for the different thermal profiles recently proposed in the literature. Three complementary arguments arise in favor of a relatively low Fe# of the Martian mantle: (i) the lack of seismic discontinuity at around 600 km depth suggests the absence of a phase transition to ringwoodite and therefore Fe# significantly lower than ~0.27. (ii) The report of one seismic discontinuity affecting both Vp and Vs is only compatible with the olivine to wadsleyite transition, which was reported to occur for Fe# lower than 0.2. (iii) A seismic transition reported at 1000 km (i.e. ~13 GPa) is compatible with our measurements with Fe# of 0.17 (± 0.02).

**1 Introduction**

Planet Mars is one of the best-studied planetary bodies in our solar system after Earth. Mars exploration programs have so far probed the planet's climate, atmosphere, surface geology, natural resources and the possibility of life (Carr et Head 2010). However, little is known about its inner structure and nature, such as its chemical composition, mineralogy, lithology, and interior dynamics. The 2018 NASA InSight mission was designed to fill this knowledge gap by probing the deep interior of the planet with several geophysical techniques such as seismic reflections method and heat-flow measurements. Previous studies on the internal structure of Mars have constrained the radius of the core (1794 ± 65 km) and its high volatile content (>16 wt% of sulfur) (Irving et al., 2023; Konopliv et al., 2020; Mocquet et al., 1996; Rivoldini et al., 2011; Sohl, Schubert and Spohn, 2005). Prior to the InSight mission, knowledge of the Martian mantle came only from experimental petrology studies and thermodynamic models (Bertka & Fei, 1997; Smrekar et al., 2018). Since then, inverse modeling based on the limited InSight seismic observations has been used to predict the internal structure of the Martian upper mantle (Drilleau et al., 2020; Amir Khan et al., 2021a), the depth and temperature at the core-mantle boundary (CMB) (Irving et al., 2023; Stähler et al., 2021) and the structure of the crust (Knapmeyer-Endrun et al., 2021).

On the other hand, geochemical constraints obtained from remote sensing, surface measurements, and the analysis of Martian meteorites (McSween & McLennan, 2014) indicate that the Martian mantle is enriched in iron compared to that of the Earth. Its Fe number (Fe# = Fe/(Mg+Fe)) was estimated between 0.21 and 0.28 (Sanloup et al., 1999; Taylor, 2013; Wänke et al., 1994; Yoshizaki & McDonough, 2019). Yoshizaki & McDonough (2019) modeled the
Fe# of the Martian mantle to be 0.21 by avoiding the use of CI chondrites as an analog of Mars and using Martian meteorites and spacecraft observations in their calculations. Finally, Khan *et al.* (2022) modeled the composition of Mars by combining geochemical with geophysical observations, including the InSight mission, which yielded a slightly lower iron content (Fe#0.19). Altogether, the Fe# of the Martian mantle appears significantly higher than the value of 0.1 of the Earth's mantle. The iron content of the mantle is crucial because the distribution of iron between mineral phases affects the thermodynamic stability of major mineral phases along the areotherm (Martian temperature profile with depth). It also affects the elastic properties of minerals and hence the seismic structure of the Martian mantle. The (Mg+Fe)/Si ratio of the mantle is also important since it controls its mineralogy and the proportion of olivine, which in turns affects the seismic velocities contrast generated by the mineralogical transformations. Experimental petrology studies show that the iron-rich basalts observed by rovers and on meteorites can only be generated by melting of a primitive iron-rich reservoir with a (Mg+Fe)/Si ratio close to that of the terrestrial peridotitic mantle (Collinet *et al.*, 2015; McSween *et al.*, 2009).

If seismic probing of the Martian mantle enables the mapping of its interior structures, laboratory data are crucial for interpreting the field-based seismic observations and refining our knowledge of the thermochemical state of the planet. Currently, there is limited experimental data on the composition of the Martian mantle, with measurements only available up to 10.6 GPa (800 km depth) (Xu *et al.*, 2021). This range is insufficient to cover the conditions of the entire Martian mantle, as the pressure at the Martian core-mantle boundary is estimated to be around 21 GPa (at 1800-2000 km, Stähler *et al.*, 2021). Based on earlier studies of phase stabilities in Fe-rich peridotite compositions, the transition of olivine to its high-pressure polymorphs (wadsleyite and ringwoodite) should occur at pressures greater than 10.6 GPa (Katsura *et al.*, 2004). The lack of experimental data dedicated to Fe-rich peridotite compositions at pressures above 10.6 GPa is a critical limitation in relating mineralogy to seismic velocities.

To constrain how phase transformations affect seismic discontinuities in the Martian mantle, we measured the velocity profiles up to a pressure of 16 GPa of two pyrolite-type compositions with iron content and mineralogy relevant to the Martian mantle (Fe# 0.17 and Fe# 0.27). We performed two different types of experiments: ultrasonic measurements at the SOLEIL synchrotron facility (PSICHE beamline, France) and “cook-and-look” experiments at the Laboratoire Magmas et Volcans (LMV, France). Both sets of experiments targeted the same pressure range with the same starting material. In the *in situ* ultrasonic experiments, the sample mineralogy was monitored using X-ray diffraction simultaneously to the velocity measurements. The LMV experiments were specifically conducted at higher temperatures to enhance grain growth and improve the chemical analyses of all phases. The recovered samples were carefully analyzed using electron probe microanalysis (EPMA), scanning electron microscopy (SEM), energy-dispersive X-ray spectroscopy (EDS) and scanning transmission electron microscopy (STEM). Additionally, we obtained information about the aggregate texture using electron backscatter diffraction (EBSD).
# 2 Materials and Methods

## 2.1 Sample preparation

Since the (Mg+Fe)/Si ratios of the Martian mantle is expected to be close to that of pyrolite (Taylor, 2013; Wang & Takahashi, 2000; Yoshizaki & McDonough, 2019), we selected iron-rich pyrolite as a starting material (e.g., Ringwood, 1962a, 1962b). Two Fe-rich compositions were prepared with similar (Mg+Fe)/Si ratio as pyrolite ((Mg+Fe)/Si=1.4), but with Fe# of 0.17 and 0.27 (Table 1). The oxide mixture was prepared using the reagent grade oxide powder and then vitrified using an aerodynamical levitation tool equipped with laser heating and under an argon flux at Conditions Extrêmes et Matériaux: Haute Température et Irradiation (CEMHTI, France). This method allows synthesis of dry glass spheres with homogeneous composition (e.g., Auzende et al., 2011). The starting glasses were synthesized under reducing conditions to avoid the crystallization of magnetite. The bulk Fe³⁺/Fe<sub>total</sub> ratio of the starting sample is the main factor controlling the redox state of the sample in these experiments performed at relatively low temperatures and for reduced heating durations.

A major advantage of using glass as a starting material for velocity measurements is the minimal development of crystallographic preferred orientation and associated anisotropy (see section 3.3). The sample crystallization occurs upon heating after compression to the target pressure, which relaxes the stress eventually accumulated in the sample during the cold compression. Relaxing the stresses is crucial to obtain reliable elasticity data. Another advantage of using glass as a starting material is the high nucleation rate, hence low grain growth, which allows for the collection of high-quality X-ray diffraction patterns from the polycrystalline aggregate and thus provides a representative measurement of the whole sample. Mineral homogeneity also implies velocity measurements representative of the bulk sample velocities. However, the small grain size produced at moderate temperatures (1000 °C) can impact on the determination of the chemical composition and thus calculation of the density of mineral phases coexisting in the recovered samples. For this reason, we performed additional experiments (“cook-and-look”) at high temperatures (1300 °C) for 6 h, to enhance grain growth.

<table><caption>**Table 1. Starting material compositions used in this study:** Fe#0.27 and Fe#0.17. Other Martian mantle compositions proposed in previous studies are also reported. D&W: Wänke *et al.*, 1994; Taylor: Taylor, 2013; Y&McD: Yoshizaki and McDonough, 2020; B&F: Bertka and Fei, 1997; Xu (UM and MM): Xu *et al.*, 2021. We also compare the various composition to the Pyrolite (Ringwood, 1962a, 1962b).</caption><thead><tr><th>Oxydes<br>(wt%)</th><th>D & W</th><th>Taylor</th><th>Y & McD</th><th>Pyrolite</th><th>Fe#0.27<br>This study</th><th>Fe#0.17<br>This study</th><th>B & F</th><th>Xu (UM)</th><th>Xu (MM)</th></tr></thead><tbody><tr><td>SiO2</td><td>44.40</td><td>43.70</td><td>45.50</td><td>46.38</td><td>42.04</td><td>43.42</td><td>43.68</td><td>46.05</td><td>45.8</td></tr><tr><td>TiO2</td><td>0.14</td><td>0.14</td><td>0.17</td><td>0.01</td><td>0</td><td>0</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Al2O3</td><td>3.02</td><td>3.04</td><td>3.59</td><td>4.53</td><td>3.17</td><td>3.27</td><td>3.13</td><td>3.25</td><td>3.3</td></tr><tr><td>MnO</td><td>0.46</td><td>0.44</td><td>0.37</td><td>0.02</td><td>0.30</td><td>0.31</td><td>-</td><td>-</td><td>-</td></tr><tr><td>FeO</td><td>17.90</td><td>18.10</td><td>14.70</td><td>8.47</td><td>21.86</td><td>15.06</td><td>18.71</td><td>18.70</td><td>18.6</td></tr><tr><td>MgO</td><td>30.20</td><td>30.50</td><td>31.00</td><td>37.47</td><td>28.62</td><td>33.79</td><td>31.50</td><td>29.60</td><td>29.8</td></tr><tr><td>CaO</td><td>2.45</td><td>2.43</td><td>2.88</td><td>2.08</td><td>2.57</td><td>2.66</td><td>2.49</td><td>2.40</td><td>2.5</td></tr></tbody></table>
<table><tbody><tr><td>Na2O</td><td>0.50</td><td>0.53</td><td>0.59</td><td>0.03</td><td>0.00</td><td>0.00</td><td>0.50</td><td>-</td><td>-</td></tr><tr><td>K2O</td><td>0.04</td><td>0.04</td><td>0.04</td><td>0.28</td><td>0.96</td><td>0.99</td><td>-</td><td>-</td><td>-</td></tr><tr><td>P2O3</td><td>0.16</td><td>0.19</td><td>0.17</td><td>0.28</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>NiO</td><td>-</td><td>0.03</td><td>0.05</td><td>0.28</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Cr2O3</td><td>0.76</td><td>0.73</td><td>0.88</td><td>0.01</td><td>0.63</td><td>0.65</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Total</td><td>100.03</td><td>99.87</td><td>99.84</td><td>99.29</td><td>100.31</td><td>100.32</td><td>100.01</td><td>100</td><td>100</td></tr><tr><td>Fe#</td><td>0.25</td><td>0.25</td><td>0.21</td><td>0.11</td><td>0.27</td><td>0.17</td><td>0.25</td><td>0.26</td><td>0.26</td></tr><tr><td>(Mg+Fe)/Si</td><td>1.35</td><td>1.39</td><td>1.29</td><td>1.4</td><td>1.45</td><td>1.45</td><td>1.45</td><td>1.35</td><td>1.37</td></tr></tbody></table>

## 2.2 In situ X-ray measurements

The white X-ray beam at the PSICHE beamline (SOLEIL synchrotron, France) is produced by an under-vacuum wiggler. X-ray energy ranges from 20 to more than 100 keV. The parameters of the beam and of other optical features of the beamline, allowing the optimal resolution in $d_{hkl}$ distance determination and the best resolution in imaging the samples, were selected following the same protocol as for previous multi-anvil study performed on PSICHE (Andrault et al., 2018). X-ray diffraction was used to monitor the sample crystallization and determine the sample mineralogy (Figure 1). Diffraction was performed in an energy-dispersive mode using a CAESAR-type diffractometer (Wang et al., 2004) with a fixed $2\theta$ angle of $8.03^{\circ} \pm 0.03^{\circ}$ determined using a gold standard. The pressure was determined based on the pressure-volume-temperature equation of state of MgO present around the sample, with an accuracy of ~0.5 GPa (Dewaele et al., 2000). Simultaneously, we acquired radiographic images for the determination of the sample length, which is required for the quantitative determination of ultrasonic velocities at each P-T condition. Since we used a focused beam configuration, each image was recorded by scanning the press in front of the beam. The samples synthesized in SOLEIL are named hereafter "in situ samples".

We observed pyroxenes both with X-ray diffraction, essentially in the olivine stability field (Supplementary Figure S2), however they have many peaks that we cannot refine cell parameters. SEM is more efficient to observe pyroxenes, especially in transition zone paragenesis where they become less abundant.

## 2.3 In situ high-pressure ultrasonic experiments

We conducted experiments from $2.9 \pm 0.5$ to $15.9 \pm 0.5$ GPa using the 1200-ton DIA-type multi-anvil module at PSICHE beamline (Figure 2). The octahedral pressure medium was composed of $MgO+Cr_2O_3$ in a 10/4 multi-anvil configuration. A dense $Al_2O_3$ buffer rod (2.1 mm in length) was placed between one of the truncations of the tungsten carbides (WC) anvils and the cylindrical glass sample of 1.2 mm in diameter and 0.6 mm in length. The buffer rod enhances the propagation of elastic waves and provides sufficient impedance contrast to reflect ultrasonic waves at the buffer rod/sample interface. We also placed a thin (between 0.3 mm and 0.7 mm thick) alumina disc on the opposite end of the sample to maintain the sample geometry. To enhance the mechanical coupling between these parts, we mirror-polished the alumina rods,
the anvil truncation and the sample ends. The use of the alumina backing plate influenced the ultrasonic signal by inversing the sample echoes due to its acoustic impedance being the same as the alumina buffer rod. Thus, the strongest amplitude to peak on the sample echo was a minimum, instead of a maximum on the buffer-rod echo. Two Rhenium disks (13 µm thick) placed at the top and bottom of the sample served as metal markers for X-ray radiography images, which aids the accurate determination of the sample length.

The sample was positioned in a MgO sleeve serving as capsule. Diffusion of Fe into the MgO capsule could occur. However, given the very sluggish kinetics of Fe-Mg inter-diffusion in periclase, the low temperatures and low durations of our *in-situ* experiments, we do not expect any significant loss of iron into the MgO capsule. Using the interdiffusion coefficients from Demouchy *et al.* (2007) diffusion distances at 900 °C and for 1 hour reach 0.3 to 3 µm. A W<sub>95</sub>Re<sub>5</sub>-W<sub>74</sub>Re<sub>26</sub> thermocouple junction was placed at one side of the sample to monitor temperatures up to 1000 °C. Heating was generated by a 25 µm thick Re foil with windows allowing the X-ray to go through the assembly. Experimental conditions are provided in **Table 2**.

The experiments comprised one, two or six compression cycles (**Supplementary figure S1**). After the first step of compression, the second step is a heating ramp at the rate of 30 °C/min. The plastic flow of the MgO assembly induces a drop in pressure at high temperatures before the gasket and cell assembly are annealed and hardened. Once pressure has stabilized at the target temperature, the sample crystallinity is monitored *in situ* as a function of time for 1 hour using X-ray diffraction. X-ray data were treated using the GSAS software (Larson, 1990). Full sample crystallization is evidenced by stationary mineralogy deduced from X-Ray diffraction patterns. Then, the ultrasonic measurements can start. The third step is the decrease of temperature, with an increment between 100 and 200 °C, and ultrasonic measurements are performed regularly until room temperature is reached. Each ultrasonic measurement cycle takes about 3 hours. Then, pressure is increased again for another cycle of heating and ultrasonic measurements.

## 2.4 Elastic wave velocity measurements

To measure the sound wave velocities across the samples, we followed the same protocol as previous studies by Chantel *et al.* (2012), Freitas *et al.* (2019), and Manthilake *et al.* (2021). This protocol is based on the ultrasonic interferometry technique (Li, Kung, and Liebermann, 2004) using electrical signals of sine waves of 20 to 50 MHz (3-5 cycles) with peak-to-peak of 5–10 V generated by an arbitrary waveform generator (Tektronix AFG3101C). The resonant frequency of the transducer is 50 MHz for the compressional waves and 30 MHz for shear waves. The waves are then converted to primary (P) and secondary (S) waves by a 10° Y-cut LiNbO<sub>3</sub> piezoelectric transducer attached to the mirror-polished truncated corner of a WC anvil. Waves propagate through the WC anvil, alumina buffer rod (BR) and the sample, and are reflected at the interfaces between anvil-BR, BR-sample, and the sample-back piece. The reflected elastic echo waves are converted back to electrical signals by the transducer and captured by a Tektronix DPO 5140 Digital Phosphor Oscilloscope at a rate of 5×10<sup>9</sup> sample/s (**Figure 3**). Signals at 20, 25, 30, 40, and 50 MHz were recorded at each temperature step during the cooling cycles. In most signals of the compressional wave (50 MHz signals), the buffer rod and anvil echoes are mixed. For this reason, we had to change the number of periods and the frequencies to distinguish the maximum of each echo. The two-way travel time for the sound
waves propagating through the sample corresponds to the time difference between the arrivals
of the echoes from the BR-sample interface and the sample-back piece interface. This travel
time is determined using the pulse-echo overlap method (Li *et al.*, 2004). We have considered
the possibility of a bond-effect on sound wave velocity (Davies et O'connell 1977; Jackson,
Niesler, and Weidner 1981; Niesler et Jackson 1989) originating from the thin Re foil placed
between the sample and the buffer rod (the use of a metal foil is necessary to identify the sample
length using X-ray radiography images, **Figure 4**). Usually, thin (~2 µm) gold disks are used
in ultrasonic setups to avoid bond-effects (Li, Kung and Liebermann, 2004). Freitas *et al.* (2019)
calculated the bond-effect induced by Re foils with thicknesses of up to 25 µm and showed that
the time shift associated with 25 µm thick Re foil was 4.59 and 8.52 ns for Vp and Vs,
respectively. The foil echoes should thus be small, both in amplitude and duration, and might
be hidden in the buffer rod and sample echoes. Given the frequency range of the transducer,
such frequency shifts are five times smaller than the width of the oscillations observed, and
therefore difficult to resolve. We thus assume that the 13 µm thick Re disks induce a twice
smaller shift, adding an uncertainty of approximately 6% on the determination of the two-way
travel time.

The main uncertainties on the velocity originate from the estimation of sample
dimensions and errors in fitting the travel time. The size of the sample is determined from the
number of pixels across the sample on the radiographic image and the sample length uncertainty
is thus of the order of the size of a pixel, i.e. 1.3 µm (Figure 4). The error in the measurement
of the travel time is estimated to be 6% as stated above. Propagating these errors induces a
relative error of less than 7% on the final velocity. The uncertainties on pressure and
temperature induce additional errors in the velocity profiles. Overall, errors in the absolute
velocities have been estimated to be less than 10% (2σ).

## 2.5 "Cook-and-look" experiments.

Additional experiments were performed *in-house* with the same glass as starting
material, but at a higher temperature (1300 °C) than the *in situ* ultrasonic experiments (**Table
2**). Enhanced grain size enables precise determination of the chemical composition of each
phase using EPMA, and the phase proportions using SEM-EBSD, while the fine-grained *in situ*
ultrasonic experiments were analysed using STEM-EDS.

Seven "cook-and-look" experiments were conducted between 5 and 17 GPa using the
1500-ton multi-anvil press of the LMV (Figure 2). We used the classical 10/4 (10M)
assembly consisting of WC cubes with 4 mm length truncations and MgO-octahedra
with 10 mm edge. The samples were loaded in MgO capsules for MA999 and
MA1000, while the other "cook-and-look" samples were loaded in Re capsules.
Temperatures up to 1300 °C were provided by a 25 µm thick rhenium foil furnace
and monitored using W/WRe thermocouples. The pressure of *in-house* experiments was
determined based on the hydraulic oil pressure of the press calibrated using well-known phase
transitions of reference materials at high temperatures (Ito, Price and Schubert, 2007) (Table
2).

## 2.6 Analytical methods

Phase proportions of the “cook-and-look” samples were obtained by evaluating the
volume occupied by each phase, using high-resolution imaging in a SEM (Zeiss SUPRA 55VP
Gemini FEG-SEM at 2MAtech, Clermont-Ferrand) with an acceleration voltage of 15 kV and working distance of 9.7 mm. The 3D volume proportion of the phases was extrapolated from the 2D imaging of the SEM by assuming the imaging is representative of the whole cylinder that is the sample. The chemical composition of relatively large mineral grains of the “cook-and-look” samples could be measured precisely with EPMA (Cameca SxFiveTactis at Laboratoire Magmas et Volcans, Clermont-Ferrand). The voltage was 15 kV and the beam current was 15 nA. We analyzed Si, Ti, Al, Cr, Mg, Fe, Ca, Ni, Mn, and Na.

Three ultrasonic samples were characterized using a SEM coupled with a focused Xe-ion beam (FIB) (Thermo Scientific™ Helios 5 PFIB CXe, in LMV, Clermont-Ferrand), equipped with an EDS and STEM detectors. High-resolution SEM images and qualitative EDS maps were acquired in backscattered electron mode at 10 kV and 0.8 nA (Figure 5A-B). In addition, 100-200 nm-thick, electron-transparent FIB foils were extracted perpendicular to the surface of the sample to be imaged and analysed quantitatively in transmission mode (STEM imaging at 30 kV and 0.1 nA; EDS analysis at 30 kV and 3.2 nA) (Figure 5C-D). The EDS analyses of the samples were calibrated by analyzing in the same conditions a FIB-foil extracted from a coarse-grained, “cook-and-look” sample (MA999) for which the compositions of the various phases were determined with EPMA. This technique allows for chemical analysis with a few nanometers resolution and with a typical ± 17 % (1 σ) uncertainty on major elements (see Gardés et al., 2020 for more details). It is well adapted for samples with sub-micrometric grains, which cannot be analysed unambiguously using EPMA as X-ray generation volumes may exceed grain volumes. Here, STEM-EDS was used to characterize three representative synchrotron experiments (MA134, MA135 and MA156). The last five experiments (MA136, MA151, MA152, MA153 and MA154) could only be chemically analysed using SEM-EDS at Geosciences Montpellier.

Grain sizes, aspect ratios and phase proportions were extracted from the indexation of electron backscatter diffraction patterns using a scanning electron microscope equipped with electron backscatter diffraction (EBSD, CamScan X500FE CrystalProbe) at Geosciences Montpellier. We acquired EBSD maps with working distance of 25.1 mm, acceleration voltage of 20 kV, beam current of 5 or 10 nA and low vacuum conditions (4 Pa of gaseous nitrogen). Detected minerals were olivine, wadsleyite, ringwoodite, pyroxenes, spinel and garnet (Supplementary figures S7 and S9). EBSD data were recorded and pre-treated with the Oxford Instruments AZtec software, before further treatment with the MTEX matlab toolbox (Bachmann et al., 2010; Hielscher & Schaeben, 2008). Data treatment includes the removal of wild spikes and filling of non-indexed pixels when seven neighbors with identical crystallographic orientations. Grain boundaries were identified where the misorientation to the next pixel was higher than 10°.
<figure><img src="image_2.png" /><figcaption>**Figure 1.** Typical X-ray diffraction patterns. All patterns were recorded in the energy dispersive mode. They show the different mineralogical assemblages observed in our samples during the *in situ* experiments. Black arrows indicate the peaks specific to wadsleyite, which were used to differentiate it from ringwoodite. Garnet is present in every pattern except those of the glass starting material (in black). X-ray diffraction patterns specific to each typical paragenesis are shown in **Supplementary Figures S2 to S6**.</figcaption></figure>

<figure><img src="image_3.png" /><figcaption>**Figure 2:** P-T conditions explored in this study for synchrotron (circles) and “cook-and-look” experiments (squares) for each Fe contents. The dashed lines emphasize qualitatively the phase transitions between olivine and its high-pressure polymorphs.</figcaption></figure>
<figure><img src="image_4.png" /></figure>

**Figure 3: Representative ultrasonic data measurements.** The ultrasonic data set recorded during the *in situ* experiments is composed of echoes corresponding to several interfaces present within the assembly. “Anvil” is the anvil/buffer rod interface. “BR maximum” is the buffer rod/sample interface, “Sample maximum” is the sample/back piece interface, and “BP maximum” is the final interface. These signals are from experiment MA154 (8.6 GPa, 920 °C). The maxima were determined using a small modeling in order to represent the overlapping between the last two echoes due to the small size of the sample (synthetic signal below the real data set).
<figure><img src="image_5.png" /><figcaption>**Figure 4. Radiographic image.** Image recorded during the ultrasonic experiments with a resolution of 1 pixel = 1.3 µm. The darkest lines are rhenium foils that mark the interfaces between the sample and the alumina rods. The furnace is composed of a rhenium foil in which two laser cut ellipsoid holes serve as low-absorbing windows for the X-ray beam. The dark horizontal edges on top and bottom of the image represent the gap between the tungsten secondary anvils. The gap is progressively closing with increasing pressure. Image dimensions: 2000 times 827 pixels.</figcaption></figure>

<figure><img src="image_6.png" /><figcaption>**Figure 5: High-resolution scanning electron microscopy on ultrasonic experiments MA135 and MA134.** (A) SEM imaging and (B) corresponding EDS map of MA135 (Fe#0.27; 11.3 GPa; 900 °C). (C) STEM imaging and (D) corresponding EDS map of the FIB foil extracted from MA134 (Fe#0.17; 14.3 GPa; 1000 °C).</figcaption></figure>
<table><caption>**Table 2: Mineralogy, phase proportion and iron content of the run products.** For olivine, wadsleyite and ringwoodite Fe# stands for Fe/(Fe+Mg). For garnets, majorite and pyroxene, calculations of Fe# include the Ca-content [Fe#=Fe/(Fe+Mg+Ca)]. Abbreviations: Ol = olivine; Wds = wadsleyite; Rwd = ringwoodite; Opx = orthopyroxene; Cpx = clinopyroxene; Grt = garnet; Maj = majorite; (Fe,Mg)O = ferropericlase; Sti = stishovite. The crosses show the phases identified by X-ray diffraction on the course of the *in situ* experiments, but which might not be necessarily present at the end of the run (due to change of P-T conditions on the course of the experiment). Only the last paragenesis could be analysed for chemical composition. Chemical compositions and modes of phases in runs MA136, MA151, MA152, MA153 and MA154 were obtained with the SEM-EDS at Geosciences Montpellier and those in runs MA134, MA135 and MA156 were obtained with the SEM-FIB at LMV Clermont-Ferrand.</caption><thead><tr><th rowspan="2">Experiments</th><th rowspan="2">Fe#</th><th rowspan="2">T(°C)</th><th rowspan="2">P (GPa)</th><th colspan="12">Mineralogy</th></tr><tr><th>Ol</th><th>Wds</th><th>Rwd</th><th>Opx</th><th>Cpx</th><th>Grt</th><th>Maj</th><th>(Fe,Mg)O</th><th>Sti</th></tr><tr><th></th><th></th><th></th><th></th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th><th>Mode:</th><th>Fe#:</th></tr></thead><tbody><tr><td colspan="18"><strong>Cook-and-look experiments:</strong></td></tr><tr><td>MA999</td><td>0.27</td><td>900</td><td>5.0</td><td>0.54</td><td>0.28</td><td></td><td></td><td>0.23</td><td>0.221</td><td>0.08</td><td>0.15</td><td>0.15</td><td>0.37</td><td></td><td></td><td></td><td></td></tr><tr><td>MA1089 (A)</td><td>0.27</td><td>1300</td><td>7.5</td><td>0.55</td><td>0.26</td><td></td><td></td><td>0.14</td><td>0.18</td><td>0.10</td><td>0.18</td><td>0.21</td><td>0.35</td><td></td><td></td><td></td><td></td></tr><tr><td>MA1093 (A)</td><td>0.27</td><td>1300</td><td>10.0</td><td>0.20</td><td>0.20</td><td></td><td>0.35</td><td>0.36</td><td>0.15</td><td>0.13</td><td>0.03</td><td>0.07</td><td></td><td></td><td>0.26</td><td>0.24</td><td></td></tr><tr><td>MA1094 (A)</td><td>0.27</td><td>1300</td><td>17.0</td><td></td><td></td><td>0.09</td><td>0.20</td><td>0.37</td><td>0.44</td><td></td><td></td><td></td><td></td><td></td><td>0.52</td><td>0.20</td><td>0.02</td><td>0.51</td></tr><tr><td>MA1081</td><td>0.27</td><td>1300</td><td>17.0</td><td></td><td></td><td></td><td></td><td>0.29</td><td>0.31</td><td></td><td></td><td></td><td></td><td></td><td>0.65</td><td>0.22</td><td>0.06</td><td>0.52</td></tr><tr><td>MA1000</td><td>0.17</td><td>1000</td><td>5.0</td><td>0.58</td><td>0.18</td><td></td><td></td><td>0.24</td><td>0.15</td><td>0.13</td><td>0.11</td><td>0.04</td><td>0.28</td><td></td><td></td><td></td><td></td></tr><tr><td>MA1089 (B)</td><td>0.17</td><td>1300</td><td>7.5</td><td>0.53</td><td>0.16</td><td></td><td></td><td>0.20</td><td>0.13</td><td></td><td></td><td>0.26</td><td>0.22</td><td></td><td></td><td></td><td></td></tr><tr><td>MA1093 (B)</td><td>0.17</td><td>1300</td><td>10.0</td><td>0.52</td><td>0.15</td><td></td><td></td><td>0.12</td><td>0.10</td><td>0.02</td><td>0.05</td><td></td><td></td><td></td><td>0.33</td><td>0.21</td><td></td></tr><tr><td>MA1087</td><td>0.17</td><td>1300</td><td>13.0</td><td></td><td></td><td>0.52</td><td>0.21</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>0.48</td><td>0.20</td><td></td></tr><tr><td>MA1094 (B)</td><td>0.17</td><td>1300</td><td>17.0</td><td></td><td></td><td>0.38</td><td>0.14</td><td>0.15</td><td>0.33</td><td></td><td></td><td></td><td></td><td></td><td>0.48</td><td>0.14</td><td></td></tr><tr><td colspan="18"><strong>Synchrotron experiments:</strong></td></tr><tr><td>MA134</td><td>0.17</td><td>1000</td><td>6.0</td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>10.5</td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>12.5</td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>14.3</td><td></td><td></td><td>0.10</td><td>0.08</td><td>0.49</td><td>0.19</td><td>0.13</td><td>0.13</td><td>0.08</td><td>0.10</td><td></td><td>0.20</td><td>0.25</td><td></td></tr><tr><td></td><td></td><td></td><td>4.9</td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>9.7</td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td></tr><tr><td>MA136</td><td>0.17</td><td>1000</td><td>10.8</td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>11.3</td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>12.4</td><td></td><td></td><td>x</td><td>x</td><td></td><td></td><td></td><td></td><td></td><td></td><td>x</td><td>x</td><td></td><td></td></tr></tbody></table>
<table><thead><tr><th rowspan="2">Experiments</th><th rowspan="2">Fe#</th><th rowspan="2">T(°C)</th><th rowspan="2">P (GPa)</th><th colspan="12">Mineralogy</th></tr><tr><th>Ol</th><th>Wds</th><th>Rwd</th><th>Opx</th><th>Cpx</th><th>Grt</th><th>Maj</th><th>(Fe,Mg)O</th><th>Sti</th></tr><tr><th></th><th></th><th></th><th></th><th>Mode:</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th><th>Mode :</th><th>Fe#:</th></tr></thead><tbody><tr><td></td><td></td><td></td><td>14.2</td><td></td><td>0.51</td><td>0.15</td><td></td><td>0.14</td><td>0.17</td><td>0.05</td><td>0.19</td><td></td><td>0.30</td><td>0.20</td><td></td><td></td><td></td></tr><tr><td>MA151</td><td>0.17</td><td>1000</td><td>10.1</td><td></td><td>0.38</td><td>0.18</td><td></td><td>0.12</td><td>0.12</td><td>0.10</td><td>0.13</td><td>0.40</td><td>0.17</td><td></td><td></td><td></td><td></td></tr><tr><td>MA152</td><td>0.17</td><td>1100</td><td>11.5</td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>×</td><td>×</td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>13.5</td><td></td><td>0.72</td><td>0.19</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>0.28</td><td>0.19</td><td></td><td></td></tr><tr><td>MA153</td><td>0.17</td><td>1000</td><td>13.3</td><td></td><td>0.32</td><td>0.19</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>0.68</td><td>0.20</td><td></td><td></td></tr><tr><td>MA135</td><td>0.27</td><td>900</td><td>3.0</td><td>×</td><td>×</td><td></td><td></td><td></td><td></td><td></td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>7.9</td><td></td><td></td><td>×</td><td>×</td><td>×</td><td>×</td><td>×</td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>9.7</td><td></td><td></td><td>×</td><td>×</td><td>×</td><td>×</td><td>×</td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>11.3</td><td></td><td>0.14</td><td>0.17</td><td>0.44</td><td>0.33</td><td>0.14</td><td>0.21</td><td>0.04</td><td>0.16</td><td></td><td>0.24</td><td>0.39</td><td></td><td></td></tr><tr><td>MA154</td><td>0.27</td><td>920</td><td>8.6</td><td>×</td><td>×</td><td></td><td></td><td>×</td><td>×</td><td></td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td></tr><tr><td></td><td></td><td>1000</td><td>10.4</td><td></td><td>0.30</td><td>0.25</td><td>0.32</td><td>0.30</td><td></td><td></td><td></td><td></td><td></td><td>0.38</td><td>0.20</td><td></td><td></td></tr><tr><td>MA156</td><td>0.27</td><td>900</td><td>14.8</td><td></td><td></td><td></td><td></td><td>×</td><td>×</td><td></td><td></td><td></td><td></td><td>×</td><td>×</td><td></td><td></td></tr><tr><td></td><td></td><td></td><td>16.0</td><td></td><td></td><td></td><td>0.57</td><td>0.29</td><td></td><td></td><td></td><td></td><td></td><td>0.42</td><td>0.25</td><td>0.01</td><td>0</td></tr></tbody></table>
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# 3 Results

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## 3.1. Mineralogy

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In our *in situ* experiments, the Fe#0.17 composition yields a mineralogical assemblage similar to that of the Earth's upper mantle. At 5 GPa, it presents $58.4 \pm 0.5$ vol.% olivine, $23.9 \pm 0.6$ vol.% orthopyroxenes, $14.3 \pm 0.7$ vol.% clinopyroxene and $4.3 \pm 0.5$ vol.% garnet (**Table 2 and Figure 6B**). With increasing pressure, the pyroxene transforms into garnet which progressively becomes majorite. At $12 \pm 1$ GPa, olivine transforms into wadsleyite ($52.4$ vol.%). At $17 \pm 1$ GPa, the mineralogy is composed of majorite ($47.6$ vol.%), wadsleyite ($37.9$ vol.%), and ringwoodite ($14.5$ vol.%) (**Table 2 and Figure 6B**). Although we did not perform experiments at pressures higher than 17 GPa, we expect the full transformation of wadsleyite into ringwoodite above 20 GPa in pyrolite compositions (Stixrude and Lithgow-Bertelloni, 2011).

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In the Fe#0.27 system, we identified four different mineral parageneses forming successively with increasing pressure. At 5 GPa, the sample is composed of $53.7 \pm 0.5$ vol.% of olivine, $22.7 \pm 0.6$ vol.% orthopyroxene, $8.3 \pm 0.7$ vol.% clinopyroxene and $15.3 \pm 0.5$ vol.% garnet (**Table 2 and Figure 6C**). With increasing pressure, the enstatite and clinopyroxene contents decrease in favor of majoritic garnet. Around 7-8 GPa, we find the co-existence of ringwoodite in equilibrium with olivine. At $12 \pm 1$ GPa, olivine transforms into a mixture of wadsleyite and ringwoodite. With further pressure increase, the amount of wadsleyite decreases in favor of ringwoodite, and above 17 GPa, the transformation into ringwoodite is complete.

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At 17 GPa, we observe traces of ferropericlase in the “cook-and-look” experiments performed up to 1300 °C, as a result of the destabilization of ringwoodite into majorite following the equation $(Mg,Fe)_2SiO_4 = (Mg,Fe)O + (Mg,Fe)SiO_3$ (**Table 2, Figures 6C and 7B**). Ferropericlase is expected according to Perple_X software, however at much higher temperatures, around 2000 °C (**Supplementary figure S14C**). Ferropericlase was not detected in the XRD spectra of the in-situ experiments. Instead, sample MA156 synthesized at 16 GPa and 900 °C in the *in-situ* experiment contains traces of stishovite (**Supplementary figure S8**). The difference is mainly due to temperature (Stixrude and Lithgow-Bertelloni, 2011; Tomioka et Miyahara, 2017). Numerical modeling of the mineralogy expected for our samples using the thermodynamic database of Stixrude and Lithgow-Bertelloni (2011) and the Perple_X software (Connolly, 2009) shows that stishovite or ferropericlase are indeed expected at low or very high temperatures, respectively (**supplementary figure S14**). However, the modal proportions are very low (**Table 2**) and the high-resolution images provide evidences that these accessory phases are not interconnected with each other (**Figure 7**). Therefore, we can reasonably assume that their impact on sound velocity is negligible. To test this assumption, we performed a fast calculation using Birch-Murnaghan (third order) equations of state to estimate the velocities using the mineralogy of MA1081 in **table 2** and replacing its ferropericlase content by either majorite or ringwoodite. The calculation with the original mineralogy of MA1081 yields Vp of 8.77 km/s and Vs of 4.65 km/s. When replacing the ferropericlase by ringwoodite the calculation yields Vp of 8.90 km/s and Vs of 4.76 km/s. Finally, when choosing majorite instead of ferropericlase, the calculation presents Vp of 8.88 km/s and Vs of 4.75 km/s. If it is not interconnected, 6 wt% of ferropericlase induces a velocity drop of -1.5% on Vp and -2.4% on Vs.

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## 3.2. Chemical composition
The “cook-and-look” samples allow to constrain the Fe content in the different phases as a function of pressure. The Fe# of olivine decreases from 0.28 at 5 GPa to 0.26 at 7.5 GPa in the Fe#0.27 system, and from 0.18 at 5 GPa to 0.15 at 7.5 GPa in the Fe#0.17 system (Table 2 and supplementary Tables S1 to S3). The Fe# decrease is due to an increasing amount of garnet (Table 2), which incorporates more iron compared to olivine, with increasing pressure (Irifune & Isshiki, 1998).

With increasing temperatures from 900 °C to 1300 °C, we observe a general increase of Fe# in olivine polymorphs, which is counterbalanced by a decreasing of Fe# in the coexisting majoritic garnet (Figure 6A and supplementary figures S10-S11). In the Fe#0.27 system, Fe# varies by only a few percent in olivine polymorphs (10-11 GPa: Fe#Wds=0.17 and Fe#Ring=0.32 at 900 °C compared to Fe#Wds=0.20 and Fe#Ring=0.36 at 1300 °C), while it varies more significantly in majoritic garnet (10-11 GPa: Fe#maj = 0.46 at 900 °C compared to Fe#maj=0.24 at 1300 °C). The effect is greater in the Fe#0.17 system, with Fe# increase in wadsleyite from 0.08 at 900 °C to 0.21 at 1300 °C. At the same pressures, Fe# decreases in majorite from 0.28 at 900 °C to 0.19 at 1300 °C.

The presence of majoritic garnet is confirmed by low Al contents and Si atom per formula unit (apfu) well above 3 (Supplementary figures S10-S11). We note that a change with temperature of the Fe partitioning between olivine-polymorphs and majoritic garnet is not expected to have a major impact on the elastic velocities of the mineral mixture typical of the Martian mantle. Increase in Fe# was reported to induce a slight decrease in the shear wave velocity in several minerals (Chung, 1970; Higo et al., 2006a; Isaak, 1992; Sinogeikin, Katsura, and Bass, 1998). However, because the increase in Fe# in olivine polymorphs is counterbalanced by a decrease in Fe# in majorite, the Fe redistribution between the two main mantle minerals should have only a second-order effect on elastic properties (see section 4.6).

<figure><img src="image_7.png" /><figcaption>**Figure 6:** (A) Phases found in the recovered samples for the two Martian compositions used in this study: ringwoodite (γ), wadsleyite (β) and olivine (α). Symbols: Blue diamonds and red squares correspond to Fe#0.17 and Fe#0.27 starting compositions, respectively. Grey lines are phase relations at 1200 °C from Katsura and Ito (1989) and Fei *et al.* (1991). The mismatches in boundaries of stability field between previous reports and our data could be due to differences in experimental temperatures. (B) and (C) are phase proportions calculated at 1300 °C as a function of pressure from compositions of phases in Fe#0.17 and Fe#0.27, respectively.</figcaption></figure>
<figure><img src="image_8.png" /><figcaption>Figure 7. Scanning electron microscopy (SEM) images of samples with different iron contents. (A): Sample MA1000 with starting composition Fe#0.17 synthesized at 5 GPa and 1000 °C. (B): Sample MA1094(A) with starting composition Fe#0.27 synthesized at 17 GPa and 1300 °C. (C): Sample MA999 with starting composition Fe#0.27 synthesized at 5 GPa and 900 °C. (D): Sample MA1094(B) with starting composition Fe#0.17 synthesized at 17 GPa and 1300 °C. All samples are homogeneous in texture and grain size regardless of pressure and temperature. Opx: orthopyroxene; Cpx: clinopyroxene; Grt: garnet; Ol: olivine; Wds: wadsleyite; Rwd: ringwoodite; Maj: majorite; (Fe,Mg)O: ferropericlase.</figcaption></figure>

We now calculate the distribution coefficient of iron ($Kd_{Fe-Mg}^{Phase1/Phase2}$) between the different polymorphs of olivine and garnet:

$$Kd_{Fe-Mg}^{Phase1/Phase2} = \frac{(Fe#Phase1 \times Mg#Phase2)}{(Fe#Phase2 \times Mg#Phase1)} \quad (1)$$

At low pressure, iron tends to preferentially distribute into garnet rather than olivine: $Kd_{Fe-Mg}^{Grt/Ol} = 1.94 \pm 0.08$ for Fe#0.27 (5 GPa, 900 °C) and $Kd_{Fe-Mg}^{Grt/Ol} = 2.21 \pm 0.09$ for Fe#0.17 (5 GPa, 1000 °C). Then, the distribution coefficient of iron between wadsleyite and majorite is close to 1, while that between ringwoodite and majorite is in favor of ringwoodite. The values of $Kd_{Fe-Mg}^{Grt/Rwd}$ range from $0.34 \pm 0.01$ to $0.71 \pm 0.03$ when there is no ferropericlase in the samples, in agreement with previous report (Frost, 2003). Only ferropericlase (FPer) incorporates more iron than ringwoodite: $Kd_{Fe-Mg}^{FPer/Rwd} = 3.20 \pm 0.13$ to $3.92 \pm 0.16$ for Fe#0.27 (17 GPa, 1300 °C). However, when analyzing the *in situ* samples, the iron distribution coefficients between garnet
and the olivine polymorphs obtained do not agree with the predictions of the literature (Frost, 2003) nor with those from Perple_X (Supplementary figure S12 and Tables S1 to S3). This is interpreted as the presence of some disequilibrium in the in-situ samples that underwent successive phase transitions. This means that in some cases, the experimental duration was too short for full sample re-equilibration. Thus, we only used the velocity measurements recorded at the first nominal pressure. There is one exception: for the Fe#0.27 composition, the wadsleyite and ringwoodite field is difficult to reach because its pressure range is only 3 GPa wide (Figure 2). We estimate that the possible lack of chemical equilibrium increases the uncertainty of the measured velocities by 3 %. Therefore, for these two experiments (MA135 and MA154), the uncertainties of measurements are 13 % (2σ).

## 3.3 Textural analysis

A total of 33 EBSD maps were acquired on seven different samples. Out of these 33 maps, 13 large-scale maps (>60000 µm²) free of charging and drifting artifacts were used to retrieve the microstructure characteristics. The step size ranges between 0.15 and 0.3 µm depending on the size of the analyzed area. The density of the orientation distribution function was calculated using an axisymmetric de la Vallee Poussin kernel with half-width of 10° (bandwidth of 28 in spherical harmonic coefficients). Euler angles per pixel (or grain) enables the detection of potential crystal preferred orientations (CPO) and the quantification of the textural strength using the J-index (Bunge, 1982). The J-index is a measurement of the intensity (or strength) of the mineral texture and is calculated as the integral of the square of the orientation distribution function. A J-index equal to 1 represents random crystallographic orientation in the aggregate, while a strong texture would lead to J-index >>5 (e.g. Demouchy et al., 2019). Representative pole figures are displayed in Figure 8. The results reveal no or very weak texture with a homogeneous arrangement of the interconnected minerals (olivine or wadsleyite; Figure 8). For the ultrasonic experiments, the maps present smaller grain sizes, but very similar textures to the “cook-and-look” experiments (Supplementary materials figure S7). In particular, the distribution and morphology of the minerals are similar: the presence of numerous small garnets (or majorites), large pyroxenes, and interconnected (Mg, Fe)₂SiO₄ polymorphs (olivine, wadsleyite or ringwoodite). The absence of the crystallographic preferred orientation of minerals and compositional heterogeneities indicate no texturally induced anisotropy (Figure 8). Thus the mean P- and S-wave velocities, only depend on the modal composition (density), temperature and pressure conditions (e.g., Tommasi & Vauchez, 2015). We also note the absence of shape-preferred orientation.

## 3.4 Global model for the elastic-wave velocities

The in situ monitoring of the sample mineralogy combined with the in situ measurement of elastic wave velocities provide constraints on the pressures at which mineralogical changes should occur as a function of the iron content, as well as the elastic wave velocities to be expected as a function of depth in the Martian mantle (Figure 9). The dependence of seismic velocity (V) on pressure and temperature can be modeled using the following formula:

$$V(P,T) = V_0 + \partial V/\partial P \times P + (\partial^2 V/\partial P \partial T \times P + \partial V/\partial T) \times (T-T_0), \quad (2)$$

where $V_0$, $\partial V/\partial P$, $\partial V/\partial T$, and $\partial^2 V/\partial P \partial T$ are the velocity at ambient conditions and its derivatives with pressure, temperature, and P-T cross parameters, respectively. We refine the
values of these four parameters based on adjustment of least-square fits to reduce the differences between the V(P,T) model and our velocity measurements at high pressure and temperature. We apply the same procedure to the different mineralogical domains relevant to the Martian mantle (Table 3 and Figure 9). In every specific mineralogy field, we use the measurements performed at varying pressure to refine the $\partial/\partial P$ term. However, as shown in Figure 6, even if the mineralogy doesn't change drastically, the phase proportions may evolve. This causes the necessity to have the $\partial^2/\partial P \partial T$ term in our model. The difference between our fits and the data ranges from 2% to 3%, with a few exceptions reaching up to 8%.

The model of elastic wave velocities refined for Fe#0.17 and Fe#0.27 can be completed with the data available for pyrolite with Fe#0.10 (i.e. preliminary reference Earth model - PREM) (Dziewonski & Anderson, 1981). We then compare the three different chemical models in the P-T fields corresponding to the same mineralogy for each iron content: olivine, pyroxenes and garnet. We consider that the four parameters of Eq. 2 vary linearly with Fe#: $X(Fe#) = A_x \times Fe# + B_x$ (Table 4). The global elastic wave velocity model that is refined enables modeling the seismic profiles in the Martian mantle as a function of its Fe#, once a temperature profile (areotherm) is defined.
<figure><img src="image_9.png" /></figure>

**Figure 8.** Crystal preferred orientations (CPO) in a lower hemisphere equal-area projection. Nb.gr: number of grains analyzed, providing the modal abundance of each phase from EBSD (with small difference from SEM-EDS, see **Table 2**). The CPO intensity is defined with the J-index corresponding to the absolute value of the integration of the orientation's distribution function. Left: “cook-and-look” sample MA1000, 5 GPa, Fe#0.27; Right: *in situ* sample MA153, 12 GPa, Fe#0.17. Note the very high Nb.gr for wadsleyite in sample MA153, which yields a statistical contouring even though the J-index remains extremely low (J=1.02, see Gasc *et al.* (2019) for further details).

505

506

507

508

509

510
Olivine Fe#17

<figure><img src="image_10.png" /></figure>

Wadsleyite Fe#17

<figure><img src="image_11.png" /></figure>

Olivine Fe#27

<figure><img src="image_12.png" /></figure>

Olivine + Ringwoodite Fe#27

<figure><img src="image_13.png" /></figure>

Wadsleyite + Ringwoodite Fe#27

<figure><img src="image_14.png" /></figure>

Ringwoodite Fe#27

<figure><img src="image_15.png" /></figure>
**Figure 9. Ultrasonic wave velocity data and models for Vp and Vs.** The scatter series and the associated uncertainties correspond to the raw data obtained with the *in situ* experiments. The lines correspond to the empirical models (Eq. 2 and Table 3) calculated at various temperatures and represented by a color gradient from room temperature to 1300 °C (the highest isotherms correspond to values extrapolated using our models). The different panels correspond to P-T ranges of stability of the polymorphs of olivine for Fe#0.17 or Fe#0.27 composition (Fe#0.17: olivine, wadsleyite) (Fe#0.27: olivine, olivine + ringwoodite, wadsleyite + ringwoodite, ringwoodite). Olivine models are based on experiments MA134 and MA136 for Fe#0.17, and MA135 for Fe#0.27. Wadsleyite models (or wadsleyite + ringwoodite) are based on experiments MA151, MA152 and MA153 for Fe#0.17 and MA135 and MA154 for Fe#0.27. Ringwoodite (or olivine + ringwoodite) models are based on experiment MA154 and MA156 both for Fe#0.27. The difference between our models and the data ranges from 2 % to 3 %, with a few exceptions reaching up to 8 %.
<table><caption>**Table 3.** Empirical models for elastic velocities of Fe#0.27 and Fe#0.17 pyrolitic compositions. This table reports the four parameters of Eq. 2 ($V_0$, $\partial V/\partial P$, $\partial V/\partial T$ and $\partial^2 V/\partial P\partial T$), for both $V_P$ and $V_S$, which are adjusted to the data sets available for compositions Fe#0.27, Fe#0.17 and Fe#0.10. n corresponds to the number of measurements used in the fit. The last three lines correspond to the UM1, UM2 and MM1 experimental models presented by Xu *et al.* (2021). The numbers in parentheses represent the standard uncertainty of the values.</caption><thead><tr><th rowspan="2">n</th><th rowspan="2">Fe#</th><th rowspan="2">Mineralogy</th><th rowspan="2">V<sub>P0</sub><br>(km.s<sup>-1</sup>)</th><th rowspan="2">V<sub>S0</sub><br>(km.s<sup>-1</sup>)</th><th colspan="2">∂V/∂P</th><th colspan="2">∂V/∂T</th><th colspan="2">∂<sup>2</sup>V/∂P∂T</th></tr><tr><th>Vp (km.s<sup>-1</sup>.GPa<sup>-1</sup>)</th><th>Vs (km.s<sup>-1</sup>.GPa<sup>-1</sup>)</th><th>Vp (km.s<sup>-1</sup>.K<sup>-1</sup>)</th><th>Vs (km.s<sup>-1</sup>.K<sup>-1</sup>)</th><th>Vp (km.s<sup>-1</sup>.K<sup>-1</sup>.GPa<sup>-1</sup>)</th><th>Vs (km.s<sup>-1</sup>.K<sup>-1</sup>.GPa<sup>-1</sup>)</th></tr></thead><tbody><tr><td>3</td><td>0.27</td><td>Ol</td><td>8.24(4)</td><td>4.43(8)</td><td>0</td><td>4.39 × 10<sup>-02</sup></td><td>-8.35 × 10<sup>-02</sup></td><td>-3.00 × 10<sup>-04</sup></td><td>0</td><td>-7.00 × 10<sup>-06</sup></td></tr><tr><td>1</td><td>0.27</td><td>Ol, Rwd</td><td>8.33(0)</td><td>4.43(0)</td><td>7.80 × 10<sup>-02</sup></td><td>6.50 × 10<sup>-02</sup></td><td>-4.66 × 10<sup>-04</sup></td><td>-2.95 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr><tr><td>4</td><td>0.27</td><td>Wds, Rwd</td><td>7.26(0)</td><td>4.39(0)</td><td>1.46 × 10<sup>-01</sup></td><td>4.92 × 10<sup>-02</sup></td><td>-2.36 × 10<sup>-04</sup></td><td>-1.90 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr><tr><td>2</td><td>0.27</td><td>Rwd</td><td>6.90(9)</td><td>1.45(9)</td><td>1.04 × 10<sup>-01</sup></td><td>2.16 × 10<sup>-01</sup></td><td>-4.96 × 10<sup>-04</sup></td><td>-3.82 × 10<sup>-04</sup></td><td>-5.18 × 10<sup>-06</sup></td><td>0</td></tr><tr><td>5</td><td>0.17</td><td>Ol</td><td>7.87(0)</td><td>4.39(6)</td><td>6.83 × 10<sup>-02</sup></td><td>4.84 × 10<sup>-02</sup></td><td>3.32 × 10<sup>-05</sup></td><td>-8.00 × 10<sup>-05</sup></td><td>0</td><td>0</td></tr><tr><td>4</td><td>0.17</td><td>Wds</td><td>8.81(4)</td><td>4.85(0)</td><td>4.00 × 10<sup>-02</sup></td><td>5.97 × 10<sup>-02</sup></td><td>-2.50 × 10<sup>-05</sup></td><td>-1.45 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr><tr><td></td><td>0.25</td><td>UM1</td><td>7.94(1)</td><td>4.53(1)</td><td>8.03 × 10<sup>-02</sup></td><td>1.72 × 10<sup>-02</sup></td><td>4.52 × 10<sup>-04</sup></td><td>2.41 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr><tr><td></td><td>0.25</td><td>UM2</td><td>7.89(1)</td><td>4.50(1)</td><td>8.30 × 10<sup>-02</sup></td><td>1.72 × 10<sup>-02</sup></td><td>4.81 × 10<sup>-04</sup></td><td>2.41 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr><tr><td></td><td>0.25</td><td>MM1</td><td>8.01(1)</td><td>4.54(2)</td><td>9.85 × 10<sup>-02</sup></td><td>3.93 × 10<sup>-02</sup></td><td>5.62 × 10<sup>-04</sup></td><td>3.72 × 10<sup>-04</sup></td><td>0</td><td>0</td></tr></tbody></table>

<table><caption>**Table 4.** Effect of iron on elastic wave velocities at high pressures and temperatures in the olivine mineralogy field. We recall as an example, the value of parameter $\partial V_P/\partial P$ of a mantle with Fe#=0.3 is equal to $\partial V_P/\partial P$ (0.3) = $A_{dVP/dP} \times 0.3 + B_{dVP/dP}$; $\partial V_P/\partial P$ (0.3) = $-6.932 \times 10^{-0.1} \times 0.3 + 1.845 \times 10^{-1}$; $\partial V_P/\partial P$ (0.3) = $-0.02346$.</caption><thead><tr><th>Parameter</th><th>A/Fe#</th><th>B</th></tr></thead><tbody><tr><td>V<sub>P0</sub> (km.s<sup>-1</sup>)</td><td>3.737</td><td>7.235</td></tr><tr><td>V<sub>S0</sub> (km.s<sup>-1</sup>)</td><td>0.422</td><td>4.324</td></tr><tr><td>∂V<sub>P</sub>/∂P (km.s<sup>-1</sup>.GPa<sup>-1</sup>)</td><td>-6.832 × 10<sup>-1</sup></td><td>1.845 × 10<sup>-1</sup></td></tr><tr><td>∂V<sub>S</sub>/∂P (km.s<sup>-1</sup>.GPa<sup>-1</sup>)</td><td>-4.452 × 10<sup>-2</sup></td><td>5.592 × 10<sup>-2</sup></td></tr><tr><td>∂V<sub>P</sub>/∂T (km.s<sup>-1</sup>.K<sup>-1</sup>)</td><td>-1.167 × 10<sup>-3</sup></td><td>2.316 × 10<sup>-4</sup></td></tr><tr><td>∂V<sub>S</sub>/∂T (km.s<sup>-1</sup>.K<sup>-1</sup>)</td><td>-2.204 × 10<sup>-3</sup></td><td>2.948 × 10<sup>-4</sup></td></tr><tr><td>∂<sup>2</sup>V<sub>P</sub>/∂T∂P (km.s<sup>-1</sup>.K<sup>-1</sup>.GPa<sup>-1</sup>)</td><td>-5.025 × 10<sup>-15</sup></td><td>-1.648 × 10<sup>-4</sup></td></tr><tr><td>∂<sup>2</sup>V<sub>S</sub>/∂T∂P (km.s<sup>-1</sup>.K<sup>-1</sup>.GPa<sup>-1</sup>)</td><td>-7.00 × 10<sup>-5</sup></td><td>1.190 × 10<sup>-5</sup></td></tr></tbody></table>
# 4 Discussion

## 4.1 Mars temperature profiles (areotherms)

The temperature profile in the Martian mantle has a crucial impact on its mineralogy and seismic structure. Higher mantle potential temperature shifts the seismic discontinuities to greater depths, due to positive Clapeyron slope of the olivine transformation to its higher-pressure polymorphs (Figure 2 and S13 in supplementary material), and decreases the seismic velocities (Eq. 2). In this work, we consider three different possible areotherms (Figure 10): the hottest is obtained assuming a liquid iron core at the core-mantle boundary (CMB) (Longhi et al., 1992). An intermediate areotherm assumes a solid iron core at the CMB (Longhi et al., 1992). The coldest profile assumes a liquid core with a Fe-FeS composition (Bertka & Fei, 1997). The last one is based on 3D models of the thermal evolution of Mars, constrained with geodetic observations and a structure coherent with a Wänke et al. (1984) composition (Plesa et al., 2018, 2021). The latest model agrees with the observations of Insight, pointing to a thick crust (between 50 and 90 km) and a large core (>1800 km of radius) (Knapmeyer-Endrun et al., 2021; Stähler et al., 2021). Plesa et al. (2018, 2021) explored the impact on the thermal profile of different amounts of heat producing elements in the lithosphere and different thicknesses of the lithosphere. As a result, their areotherms present a stronger thermal gradient in the lithosphere and a hotter mantle than in earlier studies (Figure 10). This could make the effect of temperature on seismic velocities dominant, compared to the effect of pressure in the shallow Martian mantle. We discuss this point in more detail below.

<figure><img src="image_16.png" /><figcaption>**Figure 10.** Martian thermal profiles used in this study (from Bertka & Fei, 1997; Longhi et al., 1992; Plesa et al., 2018, 2021). The adiabat from Plesa et al. (2018, 2021) is obtained by averaging all the thermal profiles reported in this previous study.</figcaption></figure>
## 4.2 Seismic velocity profiles derived from this study

Based on the elastic wave velocity models refined in this study (Eq. 2, Tables 3 and 4), we calculated the seismic profiles (Vp, Vs and Vp/Vs ratio) of the bulk Martian mantle for two possible compositions (Fe#0.17 and Fe#0.27) and the three areotherms defined above (Figure 10). We also include the depths of expected phase transitions for each given temperature profile (Figure 11 and S13 in supplementary materials), based on our experimental results (Figures 2 and 6). Focusing on the most recent areotherm (Plesa et al., 2018), the phase transitions in the Martian mantle shift to higher pressures compared to those observed in our experiments performed at lower temperatures (Figure 10). The Fe#0.27 composition displays globally lower Vp and Vs velocities than Fe#0.17 composition, in agreement with the generally observed effect of iron content in silicate minerals (Arimoto et al., 2015; Béjina et al., 2019; Higo et al., 2006b; Li and Liebermann 2000; Wang, Bass and Kastura 2014).

The present bulk-rock velocity model allows for quantifying the magnitude of the velocity jump at the phase transitions. In the Fe#0.17 system, the olivine to wadsleyite transition at 980 km depth yields velocity jumps of $\Delta$Vp of +5.42 % and $\Delta$Vs of +6.05 %, similar to those observed on Earth (PREM) (Dziewonski et Anderson, 1981). Temperature affects significantly the magnitude of the velocity jump, increasing $\Delta$Vp to +5.62% and $\Delta$Vs to +10.44% when using a colder aerotherm (compare green and red areotherms in Figure 10). In the Fe#0.27 system, at low temperatures, the first transition is the olivine to olivine + ringwoodite at 620 km depth; it is associated with $\Delta$Vp of +7.74% and $\Delta$Vs of +5.99%. Then, the olivine + ringwoodite to wadsleyite + ringwoodite transition shows $\Delta$Vp of +0.51 % and $\Delta$Vs of -0.39 % at 880 km depth. At higher temperatures, the same two transitions would be associated to velocity changes of $\Delta$Vp of +8.25 % and $\Delta$Vs of +6.72 %, and $\Delta$Vp of -2.35 % and $\Delta$Vs of -0.03 %, respectively. Due to small velocity changes, some of these phase transitions may be difficult to resolve on Mars, similarly to the wadsleyite to ringwoodite discontinuity on Earth. Still, some transitions may be better discernable using the Vp/Vs ratio, as illustrated in Figure 11. In contrast, the last transition from wadsleyite + ringwoodite to ringwoodite should be very clear: it corresponds to $\Delta$Vp of -17.88 % and $\Delta$Vs of -13.05 % at around 1130 km depth at high temperature (1750 $^{\circ}$C), or $\Delta$Vp of -13.85 % and $\Delta$Vs of -16.53 % at low temperature (1320 $^{\circ}$C).
<figure><img src="image_17.png" /></figure>

**Figure 11.** Seismic velocity models for the two compositions investigated in this study, Fe#0.17 (blue and cyan lines) and Fe#0.27 (gold and magenta lines). We calculated the seismic Vs and Vp profiles using our experimental elasticity model and the hot areotherm proposed by Plesa (2018) (cyan and gold lines) or a colder areotherm proposed by Bertka and Fei (1997) (blue and magenta lines). In orange (Fe#17) and dark blue (Fe#27) are the expected profile calculated with the Perple_X software and the elastic database of Stixrude and Lithgow-Bertelloni (2011) and the hot areotherm of Plesa (2018). We also report the seismic profile derived from the experimental data of Xu *et al.* (2021) (dashed line) who used Taylor (2013) as starting material, together with the geophysical profiles (and their uncertainties) of Khan *et al.* (2021) and Huang *et al.* (2022) (green and pink areas, respectively). The pressure and depth scales are linked to each other using a linear function (pressure ≈ depth / 74.03) using some assumptions on the structure and density of the Martian interior. The phase transitions observed in our experiments are emphasized in pink for the Fe#0.27 samples and light blue for the Fe#0.17 composition (ol: olivine, wds: wadsleyite, rwd: ringwoodite).

Interestingly, the only positive discontinuity displaying a major (i.e. easily detectable) ΔV on both Vp and Vs seismic waves in the velocity model occurs when olivine transforms to a higher-pressure polymorph (olivine to wadsleyite or olivine to ringwoodite). Since only one discontinuity with positive ΔV was inferred by InSight at a depth of 1000 km (Huang *et al.*, 2022), we can reasonably assume that it is associated with the olivine to wadsleyite transition. If ringwoodite had been involved in this transition, it would either cause a positive jump in velocities at shallower depths (< 800 km), or a major drop in velocities at higher depths (> 1100 km). This also indicates a Fe-content in the Martian mantle sufficiently low to enable the olivine to wadsleyite transition. The maximum Fe# necessary to observe this transition is not known precisely, but it should lie between our Fe#0.17 and Fe#0.27 compositions (Fei, Mao and Mysen, 1991; Katsura *et al.*, 2004; Tomioka et Miyahara, 2017).

The presence of ortho- (or clino-) pyroxenes in the Martian mantle is not expected to have a significant impact on the overall elastic wave velocities, as observed for the Earth's upper mantle. The transformation of enstatite into majoritic garnet (Kanzaki, 1987) does not show a sharp
discontinuity in density, Vp or Vs although the progressive appearance of majorite garnet may lead to a less than 5 % decrease in Vp and Vs (Arimoto *et al.*, 2015).

### 4.3 Effect of temperature profiles

An important observation is a decrease in the seismic velocities with increasing depth in the first 400-600 km for our Fe#0.27 composition, when considering the temperature profile proposed by Plesa *et al.* (2018) (Figure 11). This occurrence agrees with the previously proposed Low-Velocity Zone (LVZ) as inferred from inversion of InSight data and laboratory experiments (Khan *et al.*, 2021; Xu *et al.*, 2021). In contrast, the decrease of velocities (i.e. in the LVZ) is not observed for the Fe#0.17 composition. We investigated the seismic profiles coupled to the other possible areotherms proposed in the literature and the decrease was indeed not observed using the cold temperature profiles (Figure 11 and S13 in supplementary materials). This shows that it is the steep temperature gradient at shallow depths in the Plesa *et al.* (2018) aerotherm that induces the dominant effect of temperature over pressure on the Vp and Vs profiles in the first 400-500 km.

### 4.4 Effect of Fe# on seismic profiles

We investigated the effect of the Fe# using our global model of elastic velocities based on data available for Fe#0.1, Fe#0.17 and Fe#0.27 compositions (Table 4). Using Plesa *et al.* (2018) temperature profiles, we observe a global decrease of Vs with increasing Fe-content (Figure 12). Interestingly, we also observe a decrease of Vp and Vs with increasing depth in the shallow mantle, when the Fe# of the bulk rock is above ~0.225. Thus, the appearance of LVZ in the Martian mantle requires the conjunction of both high Fe# (>0.225) and strong temperature gradient in the lithosphere and shallow mantle (i.e. aerotherm from Plesa, 2018). Further geochemical and structural constraints are needed to discern the LVZ in the inversion of InSight's data.

### 4.5 Comparison among seismic profiles

The experimental data from Xu *et al.* (2021) present a linear model for Vp and Vs (models MM1, MM2 and UM1 in Table 3) within the olivine stability field for an analogue Fe#0.25 mantle composition of Mars. These parameters can be compared to our empirical model of Fe#0.27 pyrolite. We find a $\partial V_s/\partial P$ value 61 % higher than that proposed by Xu *et al.* (2021) for their upper mantle (UM) model (Taylor composition synthesized at 3 GPa) and only 10 % higher than their mid-mantle (MM) model (Taylor composition synthesized at 8 GPa). These differences could be related to the presence of magnetite in the sample of Xu *et al.* (2021) made at 3 GPa. Also, our Fe#0.27 profile presents a major seismic discontinuity at around 600 km depth, while the profile of Xu *et al.* (2021) for Fe#0.25 does not indicate such a feature, which could also be caused by the pressure at which their experiments were conducted (8 GPa), just below the transition. We attribute these discrepancies to the slight differences in starting materials between the two studies: according to the high-pressure phase diagram of olivine (Fei, Mao and Mysen, 1991; Tomioka and Miyahara, 2017), the lower iron content (Fe#0.25) in Xu *et al.* (2021) experiments compared to ours (Fe#0.27) is compatible with a higher transition pressure of olivine to high pressure polymorphs in their experiments.

The lack of seismic discontinuities on Vp or Vs above 800 km depth in the velocity profiles proposed by Khan (2021) suggests the absence of a phase transition from olivine to ringwoodite,
which strongly indicates significantly lower iron contents than Fe#0.27 in the Martian mantle. In the deeper parts of the Martian mantle, Huang *et al.* (2022) report a seismic discontinuity at 1000 km (~13 GPa) based on observed triplicated P and S waves. When a wavefield encounters an abrupt increase in wave speed with depth, it results in three different phases, direct, reflected and transmitted. This gives three rays with different ray parameters that emerge at the same depth. On Mars, this increase in velocities has been observed on both P and S waves around 1000 km in depth. In our study, we demonstrate that such discontinuity characteristics only happen for the olivine to wadsleyite transition. In addition, the transition pressure of 13 GPa is well compatible with the olivine to wadsleyite transition at temperatures higher than 1327 °C (Katsura *et al.*, 2004), which is compatible with the 1327 °C inferred by Huang *et al.* (2022) for Mars mantle at this depth. We can rule out a seismic discontinuity due to the olivine/ringwoodite transition at around 1000 km depths in the Martian mantle, because this transition occurs at significantly lower pressures for high Fe-contents (Fei, Mao and Mysen, 1991; Tomioka et Miyahara, 2017). Based on these observations, and with the estimation of the iron effect on the velocities (Figure 12), we constrain that the iron content in the Martian mantle ranges between Fe# of 0.17 and 0.225, depending on the aerotherm.

A final comparison of our results with seismic profiles generated by thermodynamic modeling using Perple_X coupled to the Stixrude and Lithgow-Bertelloni (2011) elastic database yields an interesting result: the Perple_X profiles for Fe#0.17 and Fe#0.27 compositions do not show any discontinuity between 6 and 12 GPa. In addition, the major drop in velocities found at high iron content in our experiments above 12 GPa is not present in the thermodynamic models which instead predicts a positive jump in velocities. Finally, the seismic velocity profiles obtained using Perple_X are less sensitive to the effect of iron, as if the mineralogy would not change significantly between Fe#0.17 and Fe#0.27. We note that the database used for iron-rich minerals in the thermodynamic modeling could be outdated (e.g. last experimental inputs for Fe-rich ringwoodite and wadsleyite are 2001 and 1998, respectively) (Stixrude and Lithgow-Bertelloni, 2011).
<figure><img src="image_18.png" /><figcaption>**Figure 12.** Seismic velocities calculated as a function of iron content. Using the parameters from **Table 4** and the temperature profile proposed by Plesa *et al.* (2018), we calculate the Vp and Vs velocity profiles and Vp/Vs ratio for four compositions, namely Fe#0.10, Fe#0.20, Fe#0.225 and Fe#0.25. It appears that increasing iron content induces a decrease of Vp and Vs, but an increase of Vp/Vs ratio.</figcaption></figure>

## 4.6 Comparison between our samples and Perple_X modelling

Comparison between the “cook-and-look” samples and the Perple_X modelling is usually very good in terms of phase abundance and phase chemistry at low pressures for both Fe#17 and Fe#27 compositions. At 1300 °C, the main differences for Fe#27 are that we find the presence of very iron-rich (Mg,Fe)O in the transition zone while Perple_X predicts the presence of stishovite. A direct consequence is that the amount of majorite is higher in the “cook-and-look” samples. Concerning the Fe#17 composition, the difference is that wadsleyite stability is wider at higher pressure: for example, we observe up to 40 % wadsleyite in sample M1094B at 17 GPa. Of course, in the “cook-and-look” samples, the pressure is less precise because we cannot measure it in-situ.

At 900 °C, the main difference is that in the *in situ* samples we find more majorite, 10 to 20 % more than in the Perple_X prediction. Also, the relative stability fields of wadsleyite and ringwoodite are different. In the Fe# in-situ samples we observe wadsleyite at 900 °C, while it is not predicted in the Perple_X modelling. Lastly, we only observe very little amount of (<1 %) stishovite at 16 GPa, while in Perple_X modelling 5 to 10 % stishovite is stable from 14 GPa.

Perple_X can predict seismic wave velocities using an elastic data base of the mineral phases. The software first estimates the elasticity (and thus elastic wave velocities) of each individual phases present in the sample at chosen P-T conditions, and uses a bounding method (Hill, 1963 a, b) to model a bulk Vp and a bulk Vs. We also calculated the bulk Vp and Vs of our samples (MA134, MA135, MA136, MA154 and MA156 in **Table 2**) using a third order Birch-Murnaghan equation of state and Voigt-Reuss-Hill bounding method (referred as BM3 hereafter) at P-T conditions identical to the experimental conditions (**Table S5**). In this case, we used the mineralogy and iron content measured on the samples and shown in **Table 2** and the elastic database is the same as that used by Perple_X (database compiled by Stixrude and Lithgow-
Bertelloni, 2011). Comparing the measured velocities to the calculated ones using BM3 provides an insight on the validity of the elastic database at high-iron content (Table S5). Then, comparing BM3 velocities to Perple_X velocities allows to quantify the effect of the mineralogy on Vp and Vs.

With the Fe#0.17 sample, we notice a 3 % difference between the Vs we measured on MA134 and MA135 and the BM3 calculation, while Vp showcases a 6 % relative difference. Such differences are well within the error bar of our measurements. As for Fe#0.27, the relative difference between our experimental results and the BM3 calculations is on average of 11 % for Vs and 12 % for Vp, with a maximum of 20 % difference for Vs with the mineralogy of MA156 (16 GPa, 900°C) and a maximum of 17 % difference for Vp with MA154's mineralogy (10.4 GPa, 900 °C). This is much higher than the experimental uncertainty and indicates that the elastic database tends to over-estimate the velocities at high iron contents.

The difference between BM3 and Perple_X velocity calculations shows less than 5 % relative differences for Fe#0.17 and Fe#0.27. An exception is that of sample MA154 (Table S5, 10.4 GPa and 900 °C), for which Perple_X predicts much lower velocities than the BM3 calculation, evidencing differences in mineralogy because both methods use the same elasticity database. Indeed, the presence of wadsleyite in the X-ray patterns of MA154 at 10.4 GPa (Figure 2) disagrees with the prediction of Perple_X showcasing olivine and ringwoodite at the same conditions (Figure S14.A).

The comparison between our experimental results and those obtained with Perple_X shows differences notably linked to: a lack of agreement on the mineralogy and divergence on the elasticity at high iron contents.

## 5 Conclusions

We report elastic wave velocity measurements of iron-enriched pyrolite aggregates obtained by a combination of multi-anvil apparatus experiments, ultrasonic interferometry and synchrotron radiation techniques. We carried out eight in situ experiments on two different bulk pyrolite compositions with Fe# of 0.17 and 0.27, over a range of pressures from 3.0 to 16.0 GPa and temperatures up to 1000 °C. These data are complemented by seven "cook-and-look" experiments between 5 and 16 GPa and temperatures varying from 900 °C to 1300 °C, using the same starting compositions. Based on these results and previous data on pyrolite composition (Fe# = 0.10), we model the effect of iron on seismic profiles along several possible Martian areotherms.

This study shows that for a high iron content (Fe# = 0.27) the Martian mantle should be composed of four layers separated by three seismic discontinuities. They would be located at ~620 km (ΔVp of +8.25 %; ΔVs of +6.72 %), ~880 km (ΔVp of +2.35 %; ΔVs of 0.03 %) and ~1130 km (ΔVp of -17.88 %; ΔVs of -13.05 %) depth. For the Fe#0.17 composition, our experiments show a single discontinuity (ΔVp of +5.62 %; ΔVs of +10.44 %) induced by the transition from olivine to wadsleyite around ~980 km of depth. Since there is no discontinuity inferred by InSight's data at depths shallower than 1000 km (Khan et al., 2021; Duràn et al., 2022; Huang et
al., 2022), we conclude that the Martian mantle should contain a lower iron content than previously estimated (i.e. significantly lower than Fe# = 0.25) (Taylor, 2013; Wänke et al., 1994). Low Fe content is compatible with previous reports of Fe# ~0.21 using cosmochemical modeling of meteorite compositions (Yoshizaki & McDonough, 2020) or combined geophysical and cosmochemical constraints (Khan et al., 2022). This could be a consequence of a depletion in FeO of the solid mantle induced by a molten lowermost mantle as reported in a recent reports (Samuel et al., 2023).

Regarding the LVZ, which was suggested at shallow depth based on the inversion of Insight's data (Khan et al., 2021), our result show that it would require a Fe content higher than 0.225 and a steep temperature gradient in the Martian lithosphere (as proposed by Plesa et al., 2018). Such a high Fe# is not compatible with the observed seismic discontinuity at 1000 km depth in the Martian mantle.

762
763 Acknowledgments

764 G.M. and N.B-C acknowledge funding from the INSU-CNRS PNP program. The Authors
765 acknowledge the synchrotron SOLEIL for the provision of beamtime under proposals 20191939,
766 20201213 and 20211446. They also thank Séverine Brassamin (CEMHTI) for assistance with the
767 starting glass preparation. The SEM-EBSD national facility at Montpellier University is supported
768 by the Institut National de Sciences de l'Univers (INSU) from Centre National de la Recherche
769 Scientifique (CNRS, France) and the Conseil Régional d'Occitanie (France). The multi-anvil
770 apparatus of Laboratoire Magmas et Volcans is financially supported by the Centre National de la
771 Recherche Scientifique (Instrument National de l'INSU). This research was financed by the French
772 Government Laboratory of Excellence initiative n°ANR-10-LABX-0006, the Région Auvergne
773 and the European Regional Development Fund. This is ClerVolc contribution number xx.

774 Author contributions

775 G.M. and N.B-C conceptualized and supervised the project. P.C. performed the experiments and
776 analysed data as part of his doctoral studies. D.A. supervised the X-ray data refinement and
777 elasticity model. L.H. and N.G. helped with the synchrotron set up at PSICHE beam-line in
778 SOLEIL Synchrotron. J.C. supervised the velocity data treatment. L. Hennet produced the glassy
779 starting materials. S.D. and F.B. performed the EDS and EBSD analysis and treatment. E.G.
780 supervised the characterization of the samples with the SEM-FIB-STEM. A.M. provided technical
781 assistance with the multi-anvil experiments. P.C. wrote the paper with inputs from all co-authors.
782 All authors participated in the discussion, read the article, and agreed on its content.

783 Competing interests: *The authors declare that they have no competing interests.*

784 Correspondence and requests for materials should be addressed to paul.chauvigne@synchrotron-soleil.fr

786 **Data and materials availability:** All the necessary data to interpret, verify and extend the
787 study is available in the manuscript, in the supplementary materials and in the following online
788 repository at: https://doi.org/10.5281/zenodo.8403132.

789

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