# Ultra-Mafic Cumulates From Kaupulehu, Hualālai Volcano, Hawai'i: Geochemical Resetting of Mantle-Inherited Olivine

A D Dujardin, Sylvie Demouchy, O Alard, Emmanuel Gardés, M Laubier, F Barou

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To cite this version:

A D Dujardin, Sylvie Demouchy, O Alard, Emmanuel Gardés, M Laubier, et al. Ultra-Mafic Cumulates From Kaupulehu, Hualālai Volcano, Hawai'i: Geochemical Resetting of Mantle-Inherited Olivine. Geochemistry, Geophysics, Geosystems, 2025, 26 (7), pp.e2024GC012128. 10.1029/2024gc012128. hal-05195276

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Distributed under a Creative Commons Attribution 4.0 International License
# Geochemistry, Geophysics, Geosystems®

## RESEARCH ARTICLE

10.1029/2024GC012128

**Key Points:**

* Olivines display no chemical zonation and have concentrations in first row transition elements between mantle and oceanic island basalt olivines

* Olivines are more enriched in trace and ultra-trace elements than typical depleted mantle-derived olivines

* Olivines display high crystallographic internal misorientation induced by dislocation creep, which is only compatible with a mantle origin

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

A. D. Dujardin,
adrianna.dujardin@uca.fr

**Citation:**

Dujardin, A. D., Demouchy, S., Alard, O., Gardés, E., Laubier, M., & Barou, F. (2025). Ultra-mafic cumulates from Kaupulehu, Hualālai volcano, Hawai'i: Geochemical resetting of mantle-inherited olivine. *Geochemistry, Geophysics, Geosystems*, 26, e2024GC012128. https://doi.org/10.1029/2024GC012128

Received 20 DEC 2024
Accepted 25 MAY 2025

**Author Contributions:**

**Conceptualization:** A. D. Dujardin,
S. Demouchy, O. Alard, E. Gardés

**Data curation:** A. D. Dujardin,
S. Demouchy

**Formal analysis:** A. D. Dujardin,
S. Demouchy, O. Alard, E. Gardés,
M. Laubier

**Funding acquisition:** S. Demouchy,
O. Alard

**Investigation:** A. D. Dujardin,
S. Demouchy, O. Alard, E. Gardés

**Methodology:** A. D. Dujardin,
S. Demouchy, O. Alard, E. Gardés,
M. Laubier, F. Barou

© 2025 The Author(s). Geochemistry, Geophysics, Geosystems published by Wiley Periodicals LLC on behalf of American Geophysical Union. This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

## Ultra-Mafic Cumulates From Kaupulehu, Hualālai Volcano, Hawai'i: Geochemical Resetting of Mantle-Inherited Olivine

<figure><img src="image_3.png" /></figure>

A. D. Dujardin¹ ID, S. Demouchy¹,² ID, O. Alard³,⁴ ID, E. Gardés¹, M. Laubier¹ ID, and F. Barou² ID

¹Laboratoire Magmas et Volcans, Université Clermont Auvergne, CNRS, IRD & OPGC, UMR, Aubière, France,

²Geosciences Montpellier, CNRS & Université de Montpellier, Montpellier, France, ³ARC Centre of Excellence for Core to Crust Fluid Systems and GEMOC National Key Centre, School of Natural Sciences, Macquarie University, Sydney, NSW, Australia, ⁴Research School of Earth Sciences, Australian National University, Canberra, ACT, Australia

**Abstract** We characterized the chemical composition and microstructure of six ultramafic cumulates from the Kaupulehu lava flow, Hualālai (Hawai'i Island), to decipher their origin. The samples are mostly wehrlites with poikilitic textures. The chemical compositions obtained from electron probe microanalyses and laser ablation inductively coupled plasma mass spectrometry confirm the magmatic origin of pyroxenes. Olivines display homogeneous compositions in major, minor and trace elements. Their Fo# (81–88) are typical of magmatic compositions, but first row transition element concentrations are intermediate between mantle and magmatic olivine compositions. Rare Earth element (REE) patterns are similar to those of mantle olivine, but Kaupulehu olivines are more enriched in heavy REE than mantle specimens. Results from Fourier transform infrared spectroscopy showed that nominally anhydrous minerals are very poor in hydrogen: 1.3–2.2 ppm H₂O by weight in olivine, 6.4 in orthopyroxene, and 12.6–48.2 in clinopyroxene. Nevertheless, fluid inclusions and bubbles evidenced by scanning electron microscopy demonstrate the presence of volatiles, which are expected to exsolve during degassing. Despite the undeniable magmatic imprint in these ultra-mafic cumulates, electron backscatter diffraction maps evidence subgrain boundaries in olivine, important internal misorientation (>10°), and Crystallographic preferred orientation (CPO) (axial [010], orthorhombic and transitional toward axial-[001]), not associated with shape preferred orientation or euhedral olivine shape. All these features are evidence of high temperature dislocation creep within the asthenospheric mantle, prior to melt percolation and chemical resetting, well before their mobilization by the volcanic eruption. We conclude that olivines are mantle-inherited, whereas clinopyroxenes result from crystallization of percolating melt.

**Plain Language Summary** We characterized the chemical composition and microstructure of six ultramafic cumulates from the Kaupulehu lava flow, Hualālai volcano (Island of Hawai'i, USA) to decipher their origin. The samples are mostly wehrlites with striking poikilitic textures, typical of magmatic cumulates. The chemical compositions attest the magmatic origin of pyroxenes, but the olivines have chemical compositions which question a mere magmatic formation. Despite the undeniable magmatic imprint in these cumulates, we used electron backscatter diffraction maps to evidence subgrain boundaries in olivine, important internal misorientation (>10°), and marked crystallographic preferred orientation (axial [010]). All these features are evidence of high temperature ductile deformation within the asthenospheric mantle, prior to melt percolation and chemical resetting, well before mobilization of the rocks by the volcanic eruption. We conclude that olivines are in fact mantle-inherited minerals, whereas pyroxenes result from crystallization of percolating melt.

## 1. Introduction

In 1800–1801, Hualālai volcano erupted and formed the Kaupulehu lava flow, which famously contains levels rich in xenoliths (i.e., nodules' bed). The most abundant xenoliths are dunites and wehrlites (60%), along with mixed mafic and ultra-mafic lithologies including dunite, wehrlite, troctolite, clinopyroxenite and gabbro (e.g., Bohrson & Clague, 1988; Chen et al., 1992; Jackson et al., 1981; Richter & Murata, 1961). Previous studies have characterized the petrology of the ultra-mafic xenoliths, especially dunites, and the trace element compositions of clinopyroxenes. These studies pointed to a magmatic origin of those cumulates rather than mantle residues from partial melting (e.g., Chen et al., 1992; White, 1966). The dunitic cumulates were interpreted as resulting from the crystallization of tholeiitic basalt evolving toward alkalic magma (Chen et al., 1992). Parental magma compositions were estimated by Gao et al. (2016, 2022). Furthermore, based on Os isotopes, recent studies have concluded that Hualālai xenoliths were the results of melt-wallrock interactions (e.g., Gao et al., 2022).
**Project administration:** S. Demouchy

**Resources:** S. Demouchy

**Supervision:** S. Demouchy, O. Alard,
E. Gardés, F. Barou

**Validation:** S. Demouchy, O. Alard,
E. Gardés, F. Barou

**Visualization:** A. D. Dujardin,
S. Demouchy

**Writing – original draft:** A. D. Dujardin,
S. Demouchy, O. Alard, E. Gardés

**Writing – review & editing:**

A. D. Dujardin, S. Demouchy, O. Alard,
E. Gardés, M. Laubier, F. Barou

The mantle underlying the Hawaiian Islands was inferred to be mineralogically and chemically heterogeneous (Clague et al., 1980) and the study of the xenolith population does not indicate the presence of a single and distinct magma reservoir below Hawai'i (Clague, 1987). These pioneer studies did not investigate trace and ultra-trace element compositions of olivine, nor did they consider the microstructure and the concentration of volatile elements, such as hydrogen, as petrogenetic indicators. The use of quantitative microstructures for ultramafics but also magmatic material is rather recent. Studies of troctolites indicated that olivine, contrary to plagioclase, was ductilely deformed in the asthenospheric mantle (Drouin et al., 2009, 2010). Other studies have used Crystallographic preferred orientation (CPO) in olivine to demonstrate ductile deformation in mantle peridotites which experienced important melt-rock interactions (e.g., Tommasi et al., 2004). A recent study has suggested that the observed internal misorientations in euhedral olivines from Kilauea are evidence of ductile deformation within crystal-mush piles (Wieser et al., 2020). The aim of this study is to constrain the petrogenesis of six representative Kaupulehu ultra-mafic xenoliths by quantifying both the trace and ultra-trace element compositions, including hydrogen, as well as the microstructures of the main mineral phases with an emphasis on olivine.

## 2. Materials and Methods

### 2.1. Geological Context and Samples

Hualālai volcano is located in the western part of Hawai'i and is one of the five volcanoes of the island (shown in Figure 1). Hualālai is about 300,000 years old, its height is 2,523 m above sea level, and it covers an area of about 850 km². The edifice is bounded to the north, east and south by Mauna Loa lava flows (Stearns & Macdonald, 1946). Hualālai, as the other volcanoes of Hawai'i, is mostly formed by tholeiitic basalt and picritic tholeiitic basalt. However, these tholeiitic flows are not exposed to the surface, as they are covered by a thin layer of late alkali basalts (Clague, 1982). The youngest eruption of Hualālai occurred in 1800–1801 from five different vents and extruded more than 300 × 10⁶ m³ of magma (Moore et al., 1987). Major and trace element compositions from whole rocks and clinopyroxenes (cpx) from Kaupulehu lava flow were reported by Chen et al. (1992). Lavas have a homogenous composition, which suggests fractionation at depth, followed by rapid magma migration toward the surface. At that time, the host alkalic basalt was estimated to derive from 5% to 10% of partial melting of a garnet-bearing mantle source enriched in light rare Earth elements (REE) (see Clague et al., 1980).

The six ultra-mafic rocks we have selected encompass two spinel-bearing dunites (NBH4, NBH7), two wehrlites with a poikilitic texture (NBH8, NBH19), one pyroxenite with poikilitic texture (NBH16), and one composite nodule with a peridotitic and a pyroxenitic section (NBH10). The major and minor element compositions of the minerals were determined using electron probe microanalyses (EPMA) and energy-dispersive X-ray spectrometry (EDS) in scanning electron microscope (SEM). The trace and ultra-trace element concentrations in the minerals were measured by laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). Hydrogen speciation and concentration were determined using Fourier-transform infrared spectrometry (FTIR). The microstructures were characterized with SEM and electron backscattered diffraction (EBSD).

### 2.2. Sample Preparation

Two sets of sections were prepared: (a) ~500 µm thick sections for FTIR and LA-ICP-MS, and (b) ~30 µm thin sections for optical microscopy, EPMA, EBSD, SEM, and EDS. Analyses by EPMA analyses were used to quantify the chemical compositions of all mineral phases, to identify potential enrichment or depletion at grain boundaries and to calculate equilibrium temperatures. EDS maps were also acquired simultaneously with EBSD maps to check for chemical homogeneity. All sections were doubly polished with diamond-lapping films with grid sizes ranging from 30 to 1/4 µm. The thick sections were prepared with crystalbond™, a thermoplastic glue which is dissoluble in acetone prior to FTIR analyses. A ~500 µm thickness has the advantage of providing sufficient material to quantify low hydrogen concentrations in the crystals (~1 wt. ppm H₂O) from FTIR in transmission mode. However, due to this large thickness, it is difficult to find a spot without cracks and defects such as inclusions and hydrous/anhydrous minerals lamellae in the beam path. For EBSD and SEM, a final polish with colloidal silica (Syton™) is necessary to remove the mechanically damaged top layer of the sample surface. Prior to EPMA and SEM, the sections were carbon-coated (a few tens of nm) but not for EBSD.
<figure><img src="image_4.png" /><figcaption>**Figure 1.** Map of Hawai'i archipelago and Hawai'i Island. The Hualālai volcano (red triangle) and its 1800–1801 Kaupulehu lava flow (in red). The vertical and horizontal axes are labeled with latitude (°N) and longitude (°W), respectively.</figcaption></figure>

## 2.3. Electron Probe Microanalyses

Major element concentrations in olivine, orthopyroxene (opx) and clinopyroxene from Kaupulehu were determined with a CAMECA SxFiveTactis electron microprobe at Laboratoire Magmas et Volcans (Université Clermont Auvergne, France) with a focused beam of 1 µm, a 15 kV accelerating voltage and a 15 nA current. Different standards were used for each element: Na from albite, Si from wollastonite or from San Carlos olivine, Ca from wollastonite, K and Al from orthose, Ti and Mn from TiMnO₃, Fe from fayalite, Mg from forsterite, and Cr from Cr₂O₃. The counting peak times ranged between 10 and 40 s depending on the element. Olivine rims and cores were analyzed to check chemical homogeneity.

## 2.4. Laser Ablation Inductively Coupled Plasma Mass Spectrometry

Trace and ultra-trace element concentrations were analyzed in olivine, orthopyroxene and clinopyroxene using an Agilent 8900 LA-ICP-MS/MS at the Department of Earth and Planetary Sciences (Macquarie University, Australia). We analyzed the cores and rims of clinopyroxene and only the cores of olivine. The ICP-MS was coupled with an automated laser ablation platform (Analyte G2) equipped with a two-volume HelEx cell and using an ArF 193 nm laser. Measurements were conducted in a He flow of about 0.8 L min⁻¹, which enhances sensitivity and reduces inter-element fractionations. The addition of a few mL min⁻¹ of H₂ increases the sensitivity for most elements by a factor of 2–4 (Guillong & Heinrich, 2007). Thus, given the expected low concentrations in trace elements of olivine, 5 mL min⁻¹ of H₂ was added to the central line (i.e., between the ablation cell and the plasma torch). The H-bearing He gas stream and particles from the sample were finally mixed with Ar (about 1 L min⁻¹) before entering the plasma with a final 30 mL mixed volume. Each analysis lasted 220 s with a 40 s pause between analyses. Analysis time comprised 140 s for the background, 60 s for mineral analysis and 20 s for wash-out monitoring. The laser was fired with an energy density of 7.5 J cm⁻², a frequency of 10 Hz, and a spot size of 110 µm. Such a large spot size is ill-suited for identifying trace element zonation.

We used NISTSRM612 (N612) glass as a calibration standard, and BCR2-G as a secondary standard. Additionally, we use MongOL as another secondary standard, which is an olivine standard developed and distributed by Batanova et al. (2019). We also analyzed SCOL and SCANU (two San Carlos olivines) as in-house long-term standards. These standards were included during analytical runs to minimize matrix effects (Bussweiler et al., 2019). Elemental concentrations were calibrated against the glass standard N612 and the natural glass BCR2-G using the values compiled by Jochum et al. (2005) in GeoRem (Table S1 in Supporting Information S1).
Data were subsequently reduced using GLITTER software (Griffin et al., 2008) by carefully inspecting the time-resolved analyses to exclude heterogeneity, such as mineral/fluid inclusions, in the analyzed volume. The SiO₂ content was used as an internal standard for reference materials and samples. Local background was used instead of the interpolated or averaged session background. This approach yielded variable detection limits for each mineral analysis, which are occasionally either lower or higher than the typical detection limit per element (best, worst, and typical detection limits are presented in Table S2 in Supporting Information S1). All elements were measured in a single acquisition. Due to the formation of polyatomic interferences during LA-ICP-MS analysis, several isotopes were measured for key elements. Polyatomic interferences can come from oxides or argides generated when oxygen or argon (even hydrogen) bind to form other atoms. For example, the formation of the Mg-argide ²⁵Mg⁴⁰Ar overlapped with ⁶⁵Cu, ¹⁴⁰Ce¹H with ¹⁴¹Pr or ¹³⁷Ba¹⁶O with ¹⁵³Eu. The mass of analyzed elements must be chosen to decrease the effect of polyatomic interference, and to keep oxide levels to a minimum (below 1%) by re-monitoring the ThO/Th ratio. ⁴⁵Sc is one of the most overestimated elements because of the formation of ¹⁶O²⁹Si. This interference can be corrected for by deconvolution. For this study, the overestimation of ⁴⁵Sc concentration is estimated to be around 2 ppm (or µg/g).

## 2.5. Fourier Transform Infrared Spectroscopy

Infrared spectra were acquired at Laboratoire Magmas et Volcans (Université Clermont Auvergne, France) with a Bruker Vertex 70 spectrometer equipped with a MCT detector (mercury cadmium telluride), which is cooled by liquid nitrogen. The MIR light was generated using a Globar™ light source and a KBr/Ge beam splitter. The spectrometer was coupled with a Hyperion-type microscope. Spectra were acquired in transmission mode and with a square aperture (100 or 200 µm) between 7,500 and 7,700 cm⁻¹ by accumulation of 300 scans with a resolution of 2 cm⁻¹. After acquisition, the spectra were corrected for background, atmospheric contamination at the crystal surface, and baseline. The three corrections were performed using OPUS software. Subsequent baseline corrections were also performed afterward with a spline function using Igor pro™ when needed. All spectra were normalized to 1 cm thickness.

Hydrogen concentration was determined using several empirical infrared calibrations: the frequency-dependent calibration of Paterson (1982) and two mineral-dependent calibrations (Withers et al., 2012 for olivine, and Bell et al., 1995 for pyroxenes). To calculate hydrogen concentration, the spectrum is integrated between 3,588 and 3,030 cm⁻¹ for olivine, between 3,700 and 3,000 cm⁻¹ for opx and between 3,690 and 3,330 cm⁻¹ for cpx.

To ease comparisons with previous studies, we still provide the hydroxyl concentration calculated using frequency-dependent calibration of Paterson (1982) as follows:

$$C_{OH} = \frac{\chi_i}{150 \xi} \int \frac{k(v)}{3780 - v} dv. \qquad (1)$$

$C_{OH}$ is the hydroxyl concentration (in wt. ppm H₂O). ξ is the orientation factor which is 1/3 for an unpolarized IR analysis (Paterson, 1982). k(v) is the absorption coefficient (in cm⁻¹), and it depends on the wavenumber v. χᵢ is the density factor which depends on chemistry. χᵢ is given by equation χᵢ = 10⁶ × 18/(2×d) where d is the density of the mineral (g/L). Typical χᵢ for mantle minerals are 2,695 wt. ppm H₂O olivine (Fo90), 2,813 wt. ppm H₂O for opx, and 2,761 wt. ppm H₂O for cpx (see Demouchy & Bolfan-Casanova, 2016 for details). The estimated error of calculated hydrogen concentrations using this calibration is around 30%.

Over the years, mineral-dependent calibrations have been an improvement compared to frequency-dependent calibrations but not without pitfalls. For olivine, the calibration of Withers et al. (2012) was selected since the calibration of Bell et al. (2003) overestimates the hydrogen concentration by 1/3. Ideally, the calibration of Withers et al. (2012) should be applied on the spectrum from each axis of a crystallographic oriented grain before summing to get the total integrated absorbance. Since the analyses were performed on thick sections and non-oriented grains, this procedure could not be applied. In this case, multiplying by three the average integrated absorbance obtained by unpolarized measurement allows to obtain a value comparable to the sum of polarized measurements acquired along the three main crystallographic axes (Férot & Bolfan-Casanova, 2012, see their Figure S1). Then, hydrogen concentration according to the calibration of Withers et al. (2012) is calculated using the formula:

$$C_{OH} = 0.119 \times \text{Abs}_{\text{tot}}, \qquad (2)$$
where $Abs_{tot}$ is the total sum of the integrated absorbance after normalization to the thickness.

To calculate the hydrogen concentration in pyroxenes, the calibration of Bell et al. (1995) is used. For unpolarized infrared measurements, this calibration also requires the total integrated absorbance

$$C_{OH} = \frac{1}{I} \times Abs_{tot}, \quad (3)$$

where I is the integral molar absorption coefficient. It is 7.09 wt. ppm H₂O cm⁻² for augite and 15.6 wt. ppm H₂O cm⁻² for enstatite (Bell et al., 1995). The detection limit was 1 wt. ppm H₂O for a sample thickness of ~500 µm. Estimated error on calculated hydrogen concentrations is around 15% for the olivine using the calibration of Withers et al. (2012) and around 20% for the pyroxenes using the calibration of Bell et al. (1995).

## 2.6. Scanning Electron Microscopy

To investigate the potential presence of hydrous minerals, melt, inclusions, bubbles along grain boundaries, and the relative chemical homogeneity of olivine (core-rim), the six samples were characterized by SEM. Back-scattered electron images were acquired at 10 kV and 0.8 nA using the SEM ThermoFisher® Helios 5 PFIB CXe at Laboratoire Magmas et Volcans (Université Clermont Auvergne, France). For the identification of phases with sizes below 10 µm, which could not be achieved by EBSD mapping, EDS was used at 10 kV and 3.2 nA. Chemical compositions were then obtained using the internal calibration of the Esprit software from Brucker™. Comparison with EPMA shows that the average uncertainty on the concentration for elements >0.3 wt. % is 8.6% relative (Figure S4 in Supporting Information S1).

## 2.7. Electron Backscatter Diffraction

The EBSD analyses were performed using a CamScan 60 CrystalProbe at Geosciences Montpellier (Université de Montpellier, France) with an acceleration voltage of 20 kV and a working distance of 25 mm. We remind that the thin sections are not carbon-coated, which permits to enhance Kikuchi band indexations. All thin sections were mapped with a step size of 10 µm, but the selected area in NBH19 was mapped with a step size of 5 µm. The electron gun was tilted at a 70° angle to the normal of the sample surface (which is here horizontal). Automatic indexations were set up for forsterite, diopside, orthoenstatite, bytownite, pargasite, chromite, and magnetite. Raw indexing rates always exceed 90%. Data were cleaned up for (a) single pixel wild orientations (spikes), (b) filling pixels which were 8, 7 and 6 indexed neighbor pixels, and (c) the pseudo symmetry which caused similar diffraction patterns. Olivine pseudosymmetry was corrected by 60° rotation around [100]. To check for chemical homogeneity, EDS maps were again acquired simultaneously with the EBSD maps (data available upon request).

Crystallographic preferred orientation data are displayed in pole figures, and presented as lower hemisphere stereographic projections. Data are presented as one measurement per grain. Thus, large crystal orientations do not create bias. Since foliation and lineation could not be identified, the orientation of the main crystallographic directions, [100], [010], and [001] for olivine and pyroxenes are plotted after a rotation into a common orientation. We impose that the maximum concentration of olivine [100] and [010] axes are rotated to be parallel to the E–W and the N–S directions of the pole figure, respectively.

The CPO calculation requires an orientation distribution function (ODF), which is calculated using a de la Vallée-Poussin kernel with a half width of 10°. The intensity of the CPO of olivine is quantified using the dimensionless J-index (Bunge, 1982), which is the volume-averaged integral of squared orientation densities. The J-index calculation is integrated into the MTEX toolbox of MatLab™ (Bachmann et al., 2011). A random CPO in crystal gives J-index equal to 1 and in naturally deformed peridotites, the J-index can reach values up to 18 (e.g., Demouchy et al., 2023). The BA-index is used to quantify the symmetry of olivine CPO (Mainprice et al., 2014). This index allows to classify the olivine CPO into three categories: (a) fiber-[010], characterized by a [010] point maximum and girdle distribution of [100] and [001] (BA-index < 0.35), (b) orthorhombic (dominant slip system [100](010)), characterized by point maxima of all [100], [010] and [001] (0.35 < BA-index < 0.65), and (c) fiber-[100], characterized by a point maxima of [100] and girdle distribution of [010] and [001] (BA-index > 0.5). The automatic grain detection routine is also obtained using the MTEX toolbox for quantitative analyses of grain sizes, aspect ratio, shape factor (grain tortuosity), and internal misorientations (using misorientation relative to mean orientation function (Mis2Mean), or the grain orientation spread function (GOS) or the grain reference
<figure><img src="image_5.png" /><figcaption>**Figure 2.** Representative optical photomicrographs in cross-polarized light of the Kaupulehu nodules. The white arrows indicate subgrain boundaries in olivine or exsolution lamellae. Note the subhedral shape of orthopyroxenes, and also the presence of dissolution/resorption features (as in antecrysts).</figcaption></figure>

orientation deviation (GROD)). Grain boundaries are defined by misorientation angles between neighboring pixels >15°, and subgrain boundaries by misorientation angles between 2 and 15°. Conventional EBSD has a detection limit of 0.5° on Euler angles. EBSD maps with a 10 µm sampling size are sufficient to determine the mineralogy and mineral fractions.

Following Tommasi et al. (2020), the internal misorientation is used as a microstructural parameter to estimate the recrystallized fraction (Rex %) of olivine, which is defined here as the area fraction of olivine grains with GOS <1.2° and grain size below 500 microns. The values are arbitrary, although they are now commonly used to estimate the fraction of recrystallized olivine in mantle-derived peridotites (e.g., Demouchy et al., 2024; Lopez-Sanchez et al., 2021; Tommasi et al., 2020).

# 3. Results

## 3.1. Mineralogy and Microstructure

The mineralogy and microstructure of the samples were investigated first using optical microscopy, and then quantitatively using EBSD maps to obtain mineral fractions and microstructure statistics, and finally using SEM to decipher the nature of the grain boundaries. Representative optical microphotographs in cross-polarized light are provided in Figure 2 (see Figure S1 in Supporting Information S1 for the entire sample sections and detailed features in Figure S2 in Supporting Information S1). The full set of EBSD phase maps and Mis2Mean maps are provided in Figure S3 in Supporting Information S1, and the most sticking microstructures are displayed in Figure 3. Representative grain boundary features are shown in Figure 4.

The lithology, mineral fractions and key microstructural parameters are compiled in Tables 1 and 2. The samples NBH4 and NBH7 are both spinel-bearing dunite with 97.5% and 99.0% olivine, and 0.9% and 0.1% spinel (chromite and magnetite), respectively. Spinels in NBH4 are euhedral or subhedral, while spinels in NBH7 are all subhedral. Rare sulfides (<0.5%) are also present. There are differences between the two dunites: NBH4 is a porphyroclastic dunite (Figure S3 in Supporting Information S1) with a bimodal grain size distribution (with peak at 200 and 2,100 µm). Note that the highest internal misorientation is found in the largest porphyroclast (up to 18°, Figure S3 in Supporting Information S1) and the parallel subgrain boundaries (Figure S3 in Supporting Information S1). On the other hand, NBH7 has an equigranular texture (Mercier & Nicolas, 1975) with a unimodal
<figure><img src="image_6.png" /><figcaption>**Figure 3.** Typical poikilitic microstructure in NBH19. (a) Phase map obtained from EBSD data showing clinopyroxene oikocrysts with olivine chadacrysts and small plagioclase patches (bytownite); (b) Euler angles map for a selected area; (c) internal crystallographic misorientation (Mis2Mean, see methods section) in olivine only showing subgrain boundaries. The recrystallized olivine grains are displayed with red grain boundaries. (d) Mis2Mean map for the selected area for olivine only in a clinopyroxene oikocryst. (e) Inverse pole figure for diopside only, showing the large grain size of clinopyroxene oikocrysts with limited internal misorientation. (f) Mis2Mean map for the selected area for a clinopyroxene patch only.</figcaption></figure>

grain size distribution (centered at 424 µm, Figure S3 in Supporting Information S1). For both samples, olivines present numerous and well-defined subgrain boundaries (Figures 2a and 2b). Such olivines are homogenously distributed at the scale of the rock section. The NBH8 sample is a spinel-bearing wehrlite with 67.3% olivine and 32.0% clinopyroxene with a poikilitic texture (Figure 2c and Figure S3 in Supporting Information S1). Again, olivines present subgrain boundaries. Clinopyroxenes commonly display exsolution lamellae of orthopyroxene and have a cuspate shape, as illustrated in Figures S2 and S3 in Supporting Information S1. The rare orthopyroxenes (0.6%) are located at the interface between olivine and clinopyroxene (Figure S3 in Supporting Information S1). Spinels are very rare and mostly euhedral. Relics of host magma with a microlithic texture are preserved at the border of this sample (Figure S1 in Supporting Information S1). The NBH10 sample (Figures S1 and S3 in Supporting Information S1) is a composite peridotite with two areas: the first one is composed almost entirely of clinopyroxene, and the second one is composed almost equally of olivine and clinopyroxene. In both areas, spinels are rare, as are orthopyroxenes (1.0%), which are only present as exsolution lamellae in clinopyroxenes or at olivine-clinopyroxene interfaces. Olivine grains display well-defined subgrain boundaries (Figure 2d). The olivine-clinopyroxene zone in NBH10 has an equigranular texture with an average grain size of around 220 µm for olivine, while the clinopyroxene-rich zone has very large grains (> mostly 1 mm). The sample NBH16 (Figure S3 in Supporting Information S1) is composed of orthopyroxenes and clinopyroxenes in almost
<figure><img src="image_7.png" /><figcaption>**Figure 4.** Backscattered electron images of grains boundaries in selected samples. Mineral phases are annotated in white and intergranular phases (glass, and bubbles) are indicated in black.</figcaption></figure>

<table><caption>**Table 1**
Lithology, Microstructure and Mineral Fractions (%) Obtained From EBSD Phase Maps (Uncertainties < 0.1%)</caption><thead><tr><th>Sample</th><th>Lithology</th><th>Microstructure</th><th>O1</th><th>Opx</th><th>Cpx</th><th>Sp</th><th>Pl</th></tr></thead><tbody><tr><td>NBH4</td><td>Dunite</td><td>Coarse-porphyroclastic</td><td>97.5</td><td>&lt;0.1</td><td>1.1</td><td>0.9</td><td>&lt;0.1</td></tr><tr><td>NBH7</td><td>Dunite</td><td>Equigranular</td><td>99.0</td><td>0</td><td>0.8</td><td>0.2</td><td>0</td></tr><tr><td>NBH8</td><td>Wehrlite</td><td>Poikilitic</td><td>67.3</td><td>0.6</td><td>31.9</td><td>&lt;0.1</td><td>&lt;0.1</td></tr><tr><td>NBH10</td><td>Composite<sup>a</sup></td><td>Equigranular</td><td>29.7</td><td>1.1</td><td>69.2</td><td>&lt;0.1</td><td>&lt;0.1</td></tr><tr><td>NBH16</td><td>Pyroxenite</td><td>Poikilitic</td><td>1.4</td><td>54.4</td><td>44.2</td><td>&lt;0.1</td><td>&lt;0.1</td></tr><tr><td>NBH19</td><td>Wehrlite</td><td>Poikilitic</td><td>85.4</td><td>&lt;0.1</td><td>12.4</td><td>&lt;0.1</td><td>2</td></tr></tbody></table>

Note. Ol, olivine; Opx, orthopyroxene; Cpx, clinopyroxene; spinel (Sp: chromite + magnetite) and plagioclase (Pl: bytownite). <sup>a</sup>The sample is composed of two equal parts (Figure S3 in Supporting Information S1), wehrlite + clinopyroxenite.
<table><caption>**Table 2**
Key Microstructural Parameters</caption><thead><tr><th rowspan="2">Sample</th><th colspan="3">Mean grain size (µm)</th><th colspan="3">Mean aspect ratio</th><th colspan="3">Mean shape factor</th><th colspan="7">Olivine</th></tr><tr><th>Ol</th><th>Opx</th><th>Cpx</th><th>Ol</th><th>Opx</th><th>Cpx</th><th>Ol</th><th>Opx</th><th>Cpx</th><th>Rex %</th><th>J-index</th><th>BA-index</th><th>Max</th><th>Mis2Mean</th><th>Mean GOS</th><th>Max GOS</th></tr></thead><tbody><tr><td>NBH4</td><td>398</td><td>44</td><td>105</td><td>1.88</td><td>1.85</td><td>2.04</td><td>1.68</td><td>1.54</td><td>1.64</td><td>25</td><td>2.30</td><td>0.64</td><td>18.5</td><td>1.13</td><td>9.6</td></tr><tr><td>NBH7</td><td>440</td><td>–</td><td>114</td><td>1.88</td><td>–</td><td>2.04</td><td>1.64</td><td>–</td><td>1.60</td><td>42</td><td>5.44</td><td>0.93</td><td>7.5</td><td>0.95</td><td>8.7</td></tr><tr><td>NBH8</td><td>378</td><td>66</td><td>311</td><td>1.85</td><td>2.4</td><td>1.91</td><td>1.65</td><td>1.76</td><td>1.82</td><td>24</td><td>1.54</td><td>0.56</td><td>11.4</td><td>1.2</td><td>6.9</td></tr><tr><td>NBH10</td><td>251</td><td>64</td><td>447</td><td>1.89</td><td>2.13</td><td>1.91</td><td>1.73</td><td>1.67</td><td>1.78</td><td>31</td><td>2.51</td><td>0.35</td><td>11.4</td><td>1.22</td><td>9.0</td></tr><tr><td>NBH16<sup>a</sup></td><td>147</td><td>96</td><td>124</td><td>1.99</td><td>2.6</td><td>2.2</td><td>1.71</td><td>1.67</td><td>1.8</td><td>n.d.</td><td>5.91</td><td>0.55</td><td>n.d</td><td>–</td><td>–</td></tr><tr><td>NBH19</td><td>309</td><td>45</td><td>157</td><td>1.76</td><td>2.12</td><td>1.90</td><td>1.62</td><td>1.61</td><td>1.76</td><td>35</td><td>1.17</td><td>0.27</td><td>11.2</td><td>1.0</td><td>10.5</td></tr></tbody></table>

Note. Mean grain sizes, aspect ratios and shape factors for ol, olivine; opx, orthopyroxene; cpx, clinopyroxene; and statistical parameters for olivine microstructure: Recrystallized fraction of olivine (Rex%), J-index (J), BA-index (BA) and angular values of maximum internal misorientation relative to the mean internal misorientation (Mis2Mean) and grain orientation spread (GOS). <sup>a</sup>Internal misorientation within opx in sample NBH16 is up to 6.4°, see Figure S3 in Supporting Information S1. n.d.: not determined.

equal proportion (54.4% and 44.2%, respectively). This spinel-bearing websterite contains very few olivines (1.4%). Both orthopyroxenes and clinopyroxenes present ubiquitous exsolution lamellae (both opx in cpx, and cpx in opx). Here, clinopyroxenes surround orthopyroxenes with a poikilitic texture and have cuspate boundaries (Figure S3 in Supporting Information S1). Several orthopyroxenes are subhedral, but many edges are often rounded and can even show evidence of dissolution/resorption (Figure S3 in Supporting Information S1). The sample NBH19 (Figure 3 and Figure S3 in Supporting Information S1) is a peridotite composed of olivine (85.4%), clinopyroxene (12.4%) and plagioclase (2%). Note that orthopyroxenes are very rare (<0.1%), and only located at olivine-clinopyroxene interfaces. Spinels are also rare (<0.1%). NBH19 has a striking poikilitic microstructure typical of magmatic cumulates with clinopyroxenes oikocrysts present as large plurimillimetric single crystals (e.g., Figure 3c, Figure S3 in Supporting Information S1), with low internal misorientation. The olivine chadacrysts are not euhedral or subhedral but rather display rounded shapes. Apart from olivine chadacrysts within clinopyroxene patches, olivine has an equigranular texture with a unimodal distribution centered on ~270 µm grain size. As for the other samples, all olivines have subgrain boundaries evidenced by optical microscopy (Figure 2) and EBSD Mis2Mean maps in Figure 3b and Figure S3 in Supporting Information S1), with internal misorientation reaching up to 11.4° locally and a GOS up to 10.5 (Table 2). Olivines with internal misorientation are homogenously distributed at the scale of the rock section (see Figure S3 in Supporting Information S1). Moreover, no shape preferred orientation (SPO) is identified (Figure 3). Minute amounts of plagioclase are present as intergranular patches (Figures 3a and 3b). On one border of the sample, relics of magma with a microlithic texture are preserved (see red arrows in Figure S2 in Supporting Information S1).

Furthermore, the detailed Euler angle map of NBH19 (white arrows in Figure 3b and Figure S3 in Supporting Information S1) shows that adjacent olivine grains inside clinopyroxene patches have the same or very similar crystallographic orientations, indicating that they are derived from the same parental grain. These grains still display internal misorientation (GOS 1–4°), but their rounded shape and orientation advocate for dissolution by a reactive melt. Nonetheless, we cannot assess whether the initial shape of the olivine grains was euhedral, subhedral, or xenomorphic.

The EBSD data allow to estimate the fraction of recrystallized olivine by spotting potentially newly formed olivine by static or dynamic recrystallization (see Section 2.7, e.g., Demouchy et al., 2024; Lopez-Sanchez et al., 2021; Tommasi et al., 2020). The fraction of recrystallized olivines ranges from 24% to 42% (Table 2). The most recrystallized sample is the small grain size dunite NBH7, and the least recrystallized sample is the composite peridotite NBH8, which has large olivine grains. Note that in many cases the rounded olivines with small but noticeable subgrain boundaries within the clinopyroxene patches are categorized as part of the recrystallized fraction of olivine, while Euler angle maps evidencing them as partially dissolved grains (Figures 3b–3d). Thus, we estimate that the calculated fraction of recrystallized olivine is overestimated.

In optical microscopy, grains show 3D defects, which could have affected the FTIR analyses. Examples of these inclusions are displayed in Figure S2 in Supporting Information S1. Three types of defects are identified: (a)
Primitive melt inclusions of 30–40 µm (Figure S2a in Supporting Information S1). These inclusions are mostly present in olivine but also in pyroxenes. They are composed of pale brown glasses with bubbles and molecular water, which affect the FTIR analyses (increase of the total absorbance up to detector saturation). (b) Clino-pyroxenes and orthopyroxenes have exsolution lamellae (Figure 2 and Figure S2 in Supporting Information S1). These exsolutions are associated with water-derived species and do not necessarily match the infrared spectra of hydrous minerals such as amphibole. (c) Array of small colorless inclusions (10 µm in average, Figures S2c and S2d in Supporting Information S1). This latter type is well-known to olivine. Their existence requires the presence of immiscible and supercritical CO₂ at high pressure during trapping of fluid inclusions (Kirby & Green, 1980; Roedder, 1965). As expected, they do not impact OH⁻ absorption IR spectra (e.g., Demouchy & Tommasi, 2021).

The microstructure of the samples was further investigated using SEM at a scale below 10 µm. Observations between grains reveal ubiquitous intergranular material, mostly crystalized as plagioclase, but we also identified glass films and bubbles. Representative images are shown in Figure 4. The composition of these phases is provided in Table 3. In all samples, the intergranular phases are distributed as films. In NBH4 and NBH7, films are a few hundred nm thick only (Figures 4a and 4b). In dunite NBH4, films are not observed systematically along all grain boundaries, while in dunite NBH7, they are always visible but represent a very low volume fraction (<<1%). The intergranular films in NBH4, NBH7 and NBH16 are too thin and could not be analyzed unambiguously by EDS. In NBH8, NBH10 and NBH16, larger films are discernible between all grains (between 1 and 5 µm thick). Films are rather rectilinear between olivine grains (Figure 4c). However, sawtooth shapes are observed at the diopside-diopside and olivine-diopside interfaces (Figures 4d and 4e). In NBH19, films are present between all grains with a thickness of a few micrometers. In several cases, the intergranular material is composed of three different phases as illustrated in Figure 4f. The main phase is plagioclase close to labradorite-bytownite compositions (Table S3 in Supporting Information S1, Deer et al., 1997). The two other phases are present as inclusions in this plagioclase. They are most probably glasses with high SiO₂ content (glass 1 and glass 2, Table S3 in Supporting Information S1). Note that glass 2 forms droplets inside glass 1 possibly indicating immiscibility during cooling. All other glass films have similar compositions (Table S3 in Supporting Information S1), with enrichment in SiO₂ ranging from 53.8 to 62.6 wt.%. The content in Al₂O₃ ranges from 17.3 to 28.9 wt.%. CaO ranges from 5.4 to 13.4 wt.%, and Na₂O ranges from 1.8 to 4.8 wt.% (see Figure S4 in Supporting Information S1 for uncertainties). These compositions do not match the composition of Hualālai lavas or parental magmas (neither tholeiitic, nor alkali basalts) previously reported by Bohrson and Clague (1988) and Gao et al. (2016, 2022). These melt compositions allow us to exclude the simple interpretation of a magma infiltration during decompression and ascent toward the surface. Overall, Kaupulehu samples are almost entirely crystallized, categorizing then as adcumulates, with <<1% of glass, which is limited to these thin films along grain boundaries (Figure 4) and overall suggest slow cooling (e.g., Mutch et al., 2019).

## 3.2. Chemical Compositions

### 3.2.1. Major Elements

Major (and selected minor) elements in olivine, orthopyroxene, clinopyroxene, spinel and plagioclase were determined using EPMA (Table 3, see also Table S4 in Supporting Information S1 for spinel and plagioclase). Key petrogenetic indicators are shown in Figure 5. Olivines have Fo# (100 × Mg/(Mg + Fe)) ranging from 80.8 (NBH-19) to 87.9 (NBH-7). Olivine nickel contents range from 0.21 to 0.45 wt.% NiO (EPMA data shown in Figure S5a in Supporting Information S1) and from 1,585 to 2,822 Ni ppm (LA-ICP-MS data, shown in Figure 5a). Clinopyroxenes are mainly augites (Table 3, Wo# < 45, Le Maitre et al., 2002) with compositions far from the range of mantle diopside compositions (Figure 5b). Spinel compositions are characterized by a Cr# (100 × Cr/(Cr + Al)), which ranges from 40.2 (NBH7) to 54.1 (NBH4). Combined with the Fo# (80.8–87.9), these compositions are well outside the olivine-spinel mantle array defined by Arai (1994). Fo# and other petrogenetic indicators in olivine are consistent with other studies on ultramafic nodules from Hualālai volcano as illustrated by Figure 5a (Bohrson & Clague, 1988; Chen et al., 1992). However, the compositions of olivines and pyroxenes of the equigranular dunite NBH7 are at the margin of the mantle-derived olivine field (Figure 5a). EPMA and EDS analyses did not show variation of major elements between olivine rim and core compositions (see Figure S6 and Table S5 in Supporting Information S1).

Equilibrium temperatures can be estimated using the partitioning of Ca between orthopyroxene and clinopyroxene from EPMA analyses following the method of Brey and Köhler (1990). Here, this geothermometer could
<table><caption>**Table 3**

Chemical Composition From EPMA for Olivine, Orthopyroxene (Opx) and Clinopyroxene (Cpx)</caption><thead><tr><th rowspan="3">wt %<br>n</th><th colspan="12">Olivine</th></tr><tr><th colspan="2">NBH4</th><th colspan="2">NBH7</th><th colspan="2">NBH8</th><th colspan="2">NBH10</th><th colspan="2">NBH16</th><th colspan="2">NBH19</th></tr><tr><th>9</th><th>1 SD</th><th>9</th><th>1 SD</th><th>7</th><th>1 SD</th><th>5</th><th>1 SD</th><th>2</th><th>1 SD</th><th>6</th><th>1 SD</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>39.76</td><td>0.14</td><td>40.33</td><td>0.23</td><td>39.84</td><td>0.21</td><td>39.581</td><td>0.090</td><td>39.58</td><td>0.15</td><td>39.15</td><td>0.20</td></tr><tr><td>TiO<sub>2</sub></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>Al<sub>2</sub>O<sub>3</sub></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>Cr<sub>2</sub>O<sub>3</sub></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>FeO</td><td>14.26</td><td>0.23</td><td>11.54</td><td>0.11</td><td>14.25</td><td>0.14</td><td>16.35</td><td>0.18</td><td>14.948</td><td>0.061</td><td>17.91</td><td>0.24</td></tr><tr><td>NiO</td><td>0.277</td><td>0.027</td><td>0.449</td><td>0.050</td><td>0.265</td><td>0.033</td><td>0.2013</td><td>0.0049</td><td>0.334</td><td>0.011</td><td>0.214</td><td>0.033</td></tr><tr><td>MgO</td><td>45.14</td><td>0.31</td><td>47.12</td><td>0.11</td><td>45.08</td><td>0.20</td><td>43.84</td><td>0.18</td><td>44.67</td><td>0.11</td><td>42.18</td><td>0.27</td></tr><tr><td>MnO</td><td>0.179</td><td>0.020</td><td>0.153</td><td>0.030</td><td>0.184</td><td>0.052</td><td>0.180</td><td>0.031</td><td>0.190</td><td>0.070</td><td>0.236</td><td>0.053</td></tr><tr><td>CaO</td><td>0.1231</td><td>0.0095</td><td>0.166</td><td>0.021</td><td>0.156</td><td>0.030</td><td>0.1281</td><td>0.0023</td><td>0.152</td><td>0.044</td><td>0.19</td><td>0.10</td></tr><tr><td>Na<sub>2</sub>O</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>K<sub>2</sub>O</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>Total</td><td>99.77</td><td></td><td>99.81</td><td></td><td>99.83</td><td></td><td>100.32</td><td></td><td>99.92</td><td></td><td>99.92</td><td></td></tr><tr><td>Fo#</td><td>84.9</td><td></td><td>87.9</td><td></td><td>84.9</td><td></td><td>82.7</td><td></td><td>84.2</td><td></td><td>80.8</td><td></td></tr><tr><th rowspan="3">n</th><th colspan="6">Opx</th><th colspan="6">Cpx</th></tr><tr><th colspan="2">NBH16</th><th colspan="2">NBH7</th><th colspan="2">NBH8</th><th colspan="2">NBH10</th><th colspan="2">NBH16</th><th colspan="2">NBH19</th></tr><tr><th>5</th><th>1 SD</th><th>3</th><th>1 SD</th><th>5</th><th>1 SD</th><th>5</th><th>1 SD</th><th>5</th><th>1 SD</th><th>5</th><th>1 SD</th></tr><tr><td>SiO<sub>2</sub></td><td>54.95</td><td>0.60</td><td>50.60</td><td>0.18</td><td>51.82</td><td>0.36</td><td>50.91</td><td>0.36</td><td>52.29</td><td>0.65</td><td>50.50</td><td>0.61</td></tr><tr><td>TiO<sub>2</sub></td><td>0.191</td><td>0.080</td><td>1.52</td><td>0.18</td><td>0.403</td><td>0.054</td><td>0.643</td><td>0.063</td><td>0.38</td><td>0.16</td><td>0.69</td><td>0.15</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>2.84</td><td>0.63</td><td>4.82</td><td>0.33</td><td>4.13</td><td>0.44</td><td>5.17</td><td>0.52</td><td>3.71</td><td>0.25</td><td>5.01</td><td>0.57</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.60</td><td>0.12</td><td>1.281</td><td>0.033</td><td>1.260</td><td>0.095</td><td>0.556</td><td>0.066</td><td>0.97</td><td>0.11</td><td>0.97</td><td>0.12</td></tr><tr><td>FeO</td><td>9.16</td><td>0.44</td><td>3.996</td><td>0.033</td><td>5.32</td><td>0.16</td><td>6.147</td><td>0.089</td><td>5.73</td><td>0.20</td><td>6.451</td><td>0.081</td></tr><tr><td>NiO</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>MgO</td><td>30.15</td><td>0.40</td><td>15.80</td><td>0.13</td><td>16.78</td><td>0.25</td><td>16.23</td><td>0.32</td><td>18.11</td><td>0.71</td><td>15.38</td><td>0.51</td></tr><tr><td>MnO</td><td>0.164</td><td>0.032</td><td>0.081</td><td>0.037</td><td>0.095</td><td>0.049</td><td>0.139</td><td>0.039</td><td>0.119</td><td>0.023</td><td>0.153</td><td>0.012</td></tr><tr><td>CaO</td><td>1.83</td><td>0.31</td><td>21.38</td><td>0.14</td><td>19.40</td><td>0.25</td><td>19.68</td><td>0.23</td><td>18.22</td><td>0.88</td><td>20.23</td><td>0.26</td></tr><tr><td>Na<sub>2</sub>O</td><td>0.063</td><td>0.013</td><td>0.613</td><td>0.051</td><td>0.58</td><td>0.11</td><td>0.475</td><td>0.034</td><td>0.488</td><td>0.036</td><td>0.537</td><td>0.031</td></tr><tr><td>K<sub>2</sub>O</td><td>0.0034</td><td>0.0049</td><td>0.0070</td><td>0.0096</td><td>0.0010</td><td>0.0014</td><td>0.0081</td><td>0.0080</td><td>0.0035</td><td>0.0033</td><td>0.0093</td><td>0.0068</td></tr><tr><td>Total</td><td>99.95</td><td></td><td>100.10</td><td></td><td>99.79</td><td></td><td>99.96</td><td></td><td>100.02</td><td></td><td>99.93</td><td></td></tr><tr><td>Mg#</td><td>85.4</td><td></td><td>87.6</td><td></td><td>84.9</td><td></td><td>82.5</td><td></td><td>84.9</td><td></td><td>81.0</td><td></td></tr><tr><td>Cr#</td><td>12.4</td><td></td><td>15.1</td><td></td><td>17.0</td><td></td><td>6.7</td><td></td><td>14.9</td><td></td><td>11.5</td><td></td></tr><tr><td>En#</td><td>82.4</td><td></td><td>47.3</td><td></td><td>49.8</td><td></td><td>48.0</td><td></td><td>52.6</td><td></td><td>45.9</td><td></td></tr><tr><td>Fs#</td><td>14.0</td><td></td><td>6.7</td><td></td><td>8.9</td><td></td><td>10.2</td><td></td><td>9.3</td><td></td><td>10.8</td><td></td></tr><tr><td>Wo#</td><td>3.6</td><td></td><td>46.0</td><td></td><td>41.4</td><td></td><td>41.8</td><td></td><td>38.1</td><td></td><td>43.3</td><td></td></tr></tbody></table>

Note. SD, standard deviation. Compositions of spinel and plagioclase are provided in Table S4 in Supporting Information S1. Fo# = 100 × Mg/(Mg + Fe); En# = (100 × Mg/(Mg + Fe + Ca)), Fs# = (100 × Fe/(Mg + Fe + Ca)) and Wo# = (100 × Ca/(Mg + Fe + Ca)).

only be applied to websterite NBH16. It yields an average equilibrium temperature of 1,180 ± 100°C. Such a high value is in agreement with the high CaO content in orthopyroxene (1.83 wt.%). However, this geo-thermobarometer was initially designed for garnet-bearing mantle-peridotites and not for spinel-bearing peridotites or magmatic cumulates (Brey & Köhler, 1990). For the orthopyroxene-free samples, we used a geothermometer based on a diffusing species slower than Fe or Mg such as aluminum in olivine (e.g., Blundy &
<figure><img src="image_8.png" /></figure>

**Figure 5.** Chemical compositions of olivine and clinopyroxene. (a) Ni as a function of Fo# in olivine. Data from previous studies are shown for comparison: Mantle olivines from various localities (Demouchy & Alard, 2021); San Carlos (Lambart et al., 2022); and French Polynesia (Tommasi et al., 2004); hydrothermal olivines from Demouchy and Alard (2021); magmatic olivines from Hawai'i and MORB (Sobolev et al., 2007); Canary Island (Neumann et al., 1999); picrite from Baffin Island (De Hoog et al., 2010); La Reunion (Demouchy & Alard, 2021); Kilauea, Hawai'i (Mourey et al., 2022); troctolite (Drouin et al., 2009). (b) TiO₂ as function Mg# in clinopyroxene for this study and previous studies Kaupulehu cumulates (Bohrson & Clague, 1988; Chen et al., 1992; Gao et al., 2016). Mantle peridotites from Mala'Ita (Ishikawa et al., 2004) and from French Polynesia (Tommasi et al., 2004).

Wood, 2003; Chakraborty, 2010). The Al-in-olivine geothermometer, measured using LA-ICP-MS, yields mean equilibrium temperatures of 988 ± 23°C and 1,084 ± 26°C using the formulations of De Hoog et al. (2010) and Bussweiler et al. (2017), respectively.

This wide range of temperature (i.e., from ≈1,200°C for pyroxenes to 1,000–1,100°C for the olivine) is similar to the temperatures previously estimated for ultramafic cumulates from Hualālai by Ziberna (2021). A pressure of 0.5 GPa was used here, following Ziberna (2021), which corresponds to <20 km of depth.

We calculated whole rock compositions using mineral compositions (Table 3) and mineral fractions (Table 1). The whole rock compositions (Figure S7 in Supporting Information S1) display linear trends which mostly reflect their mineralogy, such as a positive correlation between FeO and MgO and negative correlations between SiO₂ and MgO, Al₂O₃ and MgO, and CaO and MgO. Chin (2018) demonstrated, using fractional crystallization models, that fractional crystallization tends to produce nonlinear trends with MgO, especially for mildly incompatible elements such as Al₂O₃. Following this study, mixing of cumulates from distinct magmas could, however, explain the linear compositional trends reported here.

### 3.2.2. Minor, Trace and Ultra-Trace Elements

Minor, trace and ultra-trace concentrations in minerals were measured with LA-ICP-MS in olivine, orthopyroxene and clinopyroxene. Average concentrations in REE and Th are displayed in Figure 6 with concentrations normalized to primitive mantle (see McDonough & Sun, 1995). Extended trace element concentrations normalized to primitive mantle (McDonough & Sun, 1995) are also reported in Figure 6, and all data are provided in Table S6 in Supporting Information S1. Detection limits for olivine are provided in Table S2 of Supporting Information S1 and composition for LA-ICP-MS standards (BCR2-G and NIST612) in Table S1 of Supporting Information S1. Standard deviations are large for most of trace and ultra-trace elements in olivines (Table S6, Figures S8 and S9 in Supporting Information S1) and they reflect natural intracrystalline heterogeneities (cf. Impurities stored in point defects).

The clinopyroxene shows an enrichment in REE + Th ((Yb)N ≥ 1 (Figures 6a and 6b); where N indicates normalization to primitive mantle) with a characteristic lightly concave downward pattern showing a higher abundance of the mid-REE (MREE) relative to the heavy- and light-REE (HREE and LREE, respectively) with 1.39 < (Sm/Yb)N < 2.04 and 0.58 < (Ce/Sm)N < 0.84. The extended trace element patterns of clinopyroxene
<figure><img src="image_9.png" /><figcaption>**Figure 6.** (a, c and e) Rare Earth element patterns and (b, d and f) extended trace and ultra-trace element patterns for clinopyroxene (cpx), orthopyroxene (opx) and olivine (ol) normalized to primitive mantle (McDonough & Sun, 1995). Olivine references SCOL and SCANU are added for comparison. Black arrows indicate elements which have a concentration below the detection limit (see Table S2 in Supporting Information S1). For comparison, we added compositions for cumulative clinopyroxenes (Chen et al., 1992; Fodor & Galar, 1997; Gao et al., 2016; Sjoblom et al., 2023), pyroxenes from abyssal peridotites (Lin et al., 2023); mantle-derived orthopyroxenes (Peslier & Bizimis, 2015; Satsukawa et al., 2017; Tommasi et al., 2020); mantle-derived olivine obtained using the same analytical protocol than in this study (Demouchy & Alard, 2021) and from previous study (Satsukawa et al., 2017).</figcaption></figure>

show negative anomalies in high field strength elements (HFSE: Zr, Hf, Nb, Ta), for example, $(Nb/Ce)_N < 0.240$, and marked depletion in Cs, Rb and Ba $(\pm \text{Sr})$, (though U is not noticeably depleted relative to Th). These elements partition strongly in volatile-rich melts/fluids or in amphibole and micas, which have not been found in the NBH samples. Clinopyroxene REE patterns are typical of cumulative clinopyroxenes (e.g., Chen et al., 1992; Gao et al., 2016), and clearly distinct from those measured in residual or metasomatic mantle clinopyroxene from Hawai'i *sensus lato* (e.g., Lin et al., 2023; Sen et al., 1993; Tommasi et al., 2020). Moreover, clinopyroxene grains from ultramafic xenoliths from Mauna Kea, which were also interpreted as magmatic cumulates, have similar patterns to those from Kaupulehu (Fodor & Galar, 1997; Gao et al., 2016; Sjoblom et al., 2023; Figure 6).

The normalized REE patterns of orthopyroxene show a steadily decreasing abundance from HREE to the LREE ($(Sm/Yb)_N = 0.30$ and $(La/Sm)_N = 0.66)$ in broad equilibrium with those of clinopyroxene. This difference between LREE and HREE is also stronger when compared to mantle pyroxenes ($0.025 < (Sm/Yb)_N < 0.520$ for mantle orthopyroxenes and $0.35 < (Sm/Yb)_N < 0.55$ for mantle clinopyroxenes). Orthopyroxene REE normalized patterns of NBH16 have the same pattern shapes but are more enriched in LREE and MREE than most mantle-
derived orthopyroxenes (e.g., Peslier & Bizimis, 2015; Satsukawa et al., 2017; Tommasi et al., 2020), while their HREE concentration is lower than in orthopyroxene from abyssal peridotites (Lin et al., 2023). The extended trace element pattern shows slight positive HFSE anomalies, which are complementary to the negative ones in clinopyroxene (Figure 6d). However, the fluid mobile elements (FME: Cs, Rb, Ba ± Sr) show strong negative anomalies and a strong depletion as in the co-existing clinopyroxene.

All the olivines present a similar pattern for REE trace element with a depletion in LREE relative to HREE (Figure 6e). The shape of these patterns is often referred to as a “hockey-stick” pattern, showing a steady depletion from HREE to LREE (0.03 < (Sm/Yb)N < 0.10; 0.14 < (Ce/Sm)N < 1.10). Only La and Th show a relative enrichment. Olivines show no significant variation in REE compositions except for samples NBH7 and NBH19. The NBH7 sample is enriched in HREE and MREE, while NBH19 is depleted in LREE. The patterns of olivine inside and outside the clinopyroxene patches in NBH19 have similar shapes and concentration levels. The extended trace element diagrams for olivine show positive anomalies in HFSE and in U. Most of the FME are below detection limits in olivine. Note that NBH19 is also more depleted than other olivines for all elements except for HREE, Cs, Rb and Ba. While these olivines have patterns parallel to those measured for mantle-derived olivine analyses with the same analytical protocol (e.g., Demouchy & Alard, 2021), Kaupulehu olivines have higher concentrations in ultra-trace elements by one order of magnitude than mantle peridotites (Figures 6e and 6f). Olivine in NBH7 has trace and ultra-trace concentrations which are well above most of mantle olivines (Demouchy & Alard, 2021; Satsukawa et al., 2017). Note, however, that the highest MREE and HREE concentrations of NBH7 are similar to the most enriched olivines from Oki-dogo peridotites, which have experienced extensive melt-rock interactions at high temperatures (Satsukawa et al., 2017).

### 3.2.3. First Row Transition Elements

In magmatic cumulates, first row transition elements (Ti, V, Cr, Mn, Fe, Co, Ni, Cu, Zn and Yb) in olivine can be used as petrogenetic indicators (e.g., Foley et al., 2013; Veter et al., 2017) and selected compositions are presented in Figure 7 together with results from previous studies (e.g., De Hoog et al., 2010; Demouchy & Alard, 2021; Drouin et al., 2009; Lambart et al., 2022; Mourey et al., 2022; Neumann et al., 1999; Sobolev et al., 2007). The full data set is provided in Table S6 in Supporting Information S1 (see also Figure S5b in Supporting Information S1). Negative correlations are found between Fo# and Co, Zn and V, where Kaupulehu olivines have compositions very close to magmatic field with slight depletion in V (3–6.2 ppm) compared to previously reported magmatic compositions (6–10 ppm). Except for NBH7, most of the Kaupulehu olivines are clearly distinct from the mantle-derived olivine composition. Furthermore, the compositions in Cr and Cu help to define a specific domain for the NBH samples, which are as depleted as mantle-derived olivines but with significantly lower Fo# (Figures 7d and 7e). The concentration in Cr (66–144 ppm) is strikingly similar to that of olivines from troctolites from the mid-Atlantic ridge (Drouin et al., 2009). The concentration of Ti (39–93 ppm) combined with that of Ca (819–1,019 ppm) demonstrates that Kaupulehu olivines are more enriched in these elements than the “pristine” mantle defined by Foley et al. (2013), but less enriched in Ti and Ca than olivines from ocean island basalt (3,000–8,000 ppm; De Hoog et al., 2010; Foley et al., 2013; Neumann et al., 1999), magmatic olivines from Kilauea (Mourey et al., 2022).

### 3.2.4. Hydrogen

Representative unpolarized FTIR spectra of olivine, orthopyroxene and clinopyroxene are shown in Figure 8. All olivine spectra are characterized by an almost complete lack of hydroxyl bands except for NBH8 and NBH10 samples which show a very weak band at $3,246 cm^{-1}$. This band also appears in olivine spectra from O'ahu peridotites (Hawai'i) (Demouchy et al., 2019), but it is not attributed to a contamination by hydrous minerals or inclusion (Thoraval et al., 2019). The hydrogen concentrations in orthopyroxene could only be analyzed in NBH16 and NBH8. Only one uncontaminated spectrum could be obtained in NBH16 due to numerous inclusions and cracks in this sample. This spectrum displays only one weak band typical of orthopyroxene located at $3,602 cm^{-1}$. Orthopyroxene spectra from NBH8 show two large bands linked to contamination by other minerals and a broad band at $3,371 cm^{-1}$. For clinopyroxene in the sample NBH19, the spectrum shows a total absence of hydroxyl bands. Spectra in clinopyroxene from NBH8, NBH10 and NBH16 present the three typical OH bands of clinopyroxene located at 3,638, 3,518 and $3,450 cm^{-1}$ (Denis et al., 2015; Skogby, 2006). Clinopyroxene in samples NBH8 and NBH10 shows an additional very weak band located at $3,246 cm^{-1}$.
<figure><img src="image_10.png" /><figcaption>Figure 7.</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>**Figure 8.** Representative unpolarized FTIR spectra from orthopyroxene (blue), clinopyroxene (green) and olivine (red). All spectra are normalized to 1 cm thickness.</figcaption></figure>

Integrated absorption coefficients and corresponding hydrogen concentrations are reported in Table S7 in Supporting Information S1. Olivines which do not show the band at 3,246 cm⁻¹ (NBH4, NBH7 and NBH19) have very low hydrogen concentrations, which range between 1.3 and 2.2 wt. ppm H₂O according to the calibration of Withers et al. (2012). Olivine in NBH8 and NBH10 have a higher concentration with respectively 10.1 and 20 wt. ppm H₂O, but these higher concentrations result from the band located at 3,246 cm⁻¹ likely from an unknown

**Figure 7.** Selected first row transition elements in olivine from LA-ICP-MS. Results from this study are reported as squares (data are available in Table S6 in Supporting Information S1). Data from previous studies from different geological contexts are added for comparison: Mantle xenoliths from various localities (Demouchy & Alard, 2021); San Carlos (Lambart et al., 2022); hydrothermal olivines from Demouchy and Alard (2021); Mg-rich, magmatic olivines from Hawai'i (Sobolev et al., 2007); Canary Island (Neumann et al., 1999); picrite from Baffin Island (De Hoog et al., 2010); La Reunion (Demouchy & Alard, 2021); Kilauea, Hawai'i (Mourey et al., 2022); troctolite (Drouin et al., 2009).
<figure><img src="image_12.png" /><figcaption>**Figure 9.** Olivine, orthopyroxene and clinopyroxene crystallographic preferred orientations (lower hemisphere projection). Data are represented as density contours when the number of analyzed grains is greater than 100. When the number of grains is lower than 100, data are shown as points (see Figure S3 in Supporting Information S1 for details).</figcaption></figure>

contaminating phase. The only uncontaminated measurement of orthopyroxene yields a low hydrogen concentration of 6.4 wt. ppm H₂O. Hydrogen concentrations in clinopyroxenes range between 4.8 and 48.2 wt. ppm H₂O. These concentrations are very low for both pyroxenes when compared to the current database on nominally anhydrous minerals (Demouchy & Bolfan-Casanova, 2016). The FTIR measurements demonstrate that all minerals are hydrogen poor, which is in agreement with previous studies of mantle-derived peridotite xenoliths from O'ahu, another island of the Hawaiian archipelago (~0.8 wt. ppm of H₂O in olivine and ~19.5 wt. ppm H₂O in orthopyroxene, Demouchy & Tommasi, 2021; Peslier & Bizimis, 2015).

The determination of hydrogen concentration in individual minerals and mineral fractions (Table 1 and Figure S7 in Supporting Information S1) permits the estimation of the whole rock water content based on NAMs only. Concentrations ranges around 1.3–1.4 wt. ppm H₂O for dunites and 2.5–39.3 wt. ppm H₂O for the four other samples (see Table S7 in Supporting Information S1). Note that these concentrations do not consider water and water-derived species partitioned into intergranular phases (e.g., glasses and bubbles).

### 3.3. Crystal Preferred Orientation

Pole figures of olivine, orthopyroxene and clinopyroxene are presented in Figure 9. Pole figures of olivine from sample NBH4 show point maxima for [100], and [001] axes with a dislocation glide (point concentration) subparallel to [100], and (010) planes almost normal to the shear plane (in top insert in Figure 9), which display a girdle. This CPO attests the activation of the [100](010) slip system in olivine (also call the A-type, Jung & Karato, 2001). In sample NBH7, pole figures have point maxima in [100] direction and a well-defined girdle in [010] but a weaker girdle for [001] axis. This CPO is intermediate between A-type and D-type (Jung & Karato, 2001), typical of the activation of the [100](0 kl) slip system in olivine. Pole figures for NBH8 show a well-
more complex petrogenesis than simple magmatic crystallization and suggest a modification of mantle olivine compositions by melt-rock interactions (e.g., Bodinier & Godard, 2014; Tommasi et al., 2004).

Major and minor elements in silicates are expected to have faster mobilities than trace and ultra-trace elements in silicates when equilibrated with melts or fluids (e.g., Chakraborty, 2010). Therefore, trace and ultra-trace elements are commonly used to fine-tune chemical interactions such as cryptic metasomatic events or melt-rock interactions (e.g., O'Reilly & Griffin, 2013). REE and extended trace element patterns for Kaupulehu olivines are similar to mantle-derived olivines (hockey-stick shape). These patterns show a steady depletion from HREE to LREE except for the most incompatible REE (La ± Ce) (Demouchy & Alard, 2021; Lin et al., 2023; Satsukawa et al., 2017), but concentrations are higher in Kaupulehu olivines by one order of magnitude (Figure 6). For example, Lu concentration ranges between 3.77 and 8.31 ppb, while Lu concentration ranges from 0.25 to 3.82 ppb for mantle olivines (Demouchy & Alard, 2021). A similar comparison can be made for La (0.27 ≤ [La]OI-NBH ≤ 2.09 ppb vs. 0.09 ≤ [La]OI-Mantle ≤ 1.19 ppb, Demouchy & Alard, 2021). Note that the concentrations in HREE and MREE in NBH7 are higher than other Kaupulehu olivines and close to mantle-inherited olivines from Oki-Dogo (Figure 6), which are typical examples of extensive melt-rock interactions in a hot subduction context (Satsukawa et al., 2017). Overall, the olivine REE patterns suggest a mantle composition modified by melt-rock interactions without reaching equilibrium for trace and ultra-trace elements.

Based on the experimental measurements of ionic diffusion in olivine reported in the literature (Coogan et al., 2005; Dohmen et al., 2007; Petry et al., 2004; Spandler et al., 2007 and a review by Chakraborty, 2010), Fe-Mg, Ni, Co, Cr, Mn, Ca, Ti, V diffusivities range between $10^{-17}$ and $10^{-15}$ m²·s⁻¹ at 1,200°C. Thus, 2–200 years would be necessary to reset the chemical composition of a 1 mm grain of olivine, which is rapid enough for being possible given the volcanic activity of Hawai'i Island. As mentioned above, trace and ultra-trace elements are expected to be slower diffusing species in olivine (Chakraborty, 2010), but available diffusion coefficients for REEs are in disagreement. Spandler and O'Neill (2010) reported REE diffusivities of ~$10^{-16}$ m²·s⁻¹ at 1,300°C, while Cherniak (2010) reported diffusivities of $10^{-20}$–$10^{-19}$ m²·s⁻¹ at 1,200°C. Thus, it would take ~20 years to reset the REE compositions in a grain size of 1 mm according to the former, and ~20,000 years according to the latter. Only this second case could have led to the preservation of the current REE hockey-stick pattern of olivine.

Furthermore, most olivines in the clinopyroxene patches are rounded, which suggests a dissolution of olivine edges in the melt. The kinetics of olivine dissolution in basaltic melt can be very fast at high temperature. For instance, the reduction of an olivine with a 4 mm diameter grain to 1 mm takes only 15 days at 1,200°C according to the dissolution law of Chen and Zhang (2008) as long as the melt does not reach MgO-saturation.

All NAMs are hydrogen-poor (1.3–2.2 wt. ppm H₂O in olivine, 6.4 wt. ppm H₂O in orthopyroxene and between 4.8 and 48.2 wt. ppm H₂O in clinopyroxene) relative to rare mantle specimens from the oceanic lithosphere (Demouchy & Bolfan-Casanova, 2016; Demouchy et al., 2015, 2017, 2024). These low concentrations in hydrogen are expected and are attributed to two processes: (a) The high H incompatibility during melting (Dixon et al., 2002) with a peridotite/melt partition coefficient between $4.10^{-3}$ and $8.10^{-3}$ at upper mantle conditions (Hirschmann et al., 2009) and (b) the fast diffusivity of hydrogen in minerals and along grain boundaries (e.g., Demouchy, 2010; Demouchy & Mackwell, 2006). Hydrogen thus accumulates in the liquid phase, exsolves and forms bubbles upon decompression (Figure 4). The fast diffusivity of hydrogen in olivine implies its inadequacy as a petrogenetic indicator to infer olivine petrogenesis as H concentration will trend to re-equilibrate down to low temperatures. Therefore, we must rely on non-chemical criteria to further establish the petrogenesis of olivine.

## 4.2. Evidence of Mantle Inheritance From Microstructures

The olivine microstructures provide additional key observations. First, olivine grains are not euhedral, nor subhedral, even when they are trapped within clinopyroxene oikocrysts, unlike what would be expected for magmatic phenocrysts (Mutch et al., 2019; Wieser et al., 2020). Here, olivine edges appear dissolved by a reactive melt, which then crystallized as clinopyroxene (Figure 3, Figure S3 in Supporting Information S1). In parallel, the cuspate shape of clinopyroxenes and plagioclases, the lack of CPOs in pyroxenes and the lack of correlation with olivine CPOs indicate that they crystallized after olivine in NBH8, NBH10 and NBH19. While the resorption of magmatic olivine is possible (e.g., Bouvet de Maisonneuve et al., 2016), olivine chadacrysts are usually subhedral (Mutch et al., 2019) and not strongly deformed (Mis2Mean < 5°). Moreover, the deformed Kaupulehu olivines are homogenously distributed at the scale of the thin section and present well-developed CPOs (Figure 9). These CPOs evidence the activation of [100] dislocations in most samples ([100](0kl) slip system) except for sample
NBH19, which displays a transition between two slip systems [100](010) and [001](010) (e.g., Demouchy et al., 2023; Holtzman et al., 2003; Jung & Karato, 2001). These slip systems of olivine are typical of dislocation creep at high temperature commonly described in mantle peridotites (e.g., Demouchy et al., 2023; Tommasi et al., 2000). Olivine grains do not show SPO, while it is a common microstructural feature of magmatic minerals acquired during lava flow emplacement (e.g., Vernon, 2000).

Subgrain boundaries and well-developed CPOs in olivines from melt-rich contexts such as mid-ocean ridges are considered as strong arguments for ductile deformation acquired in the mantle (e.g., Boudier & Nicolas, 1995; Demouchy et al., 2023; Drouin et al., 2009; Drouin et al., 2010). Likewise, in intra-oceanic settings, dunites from French Polynesia (Tommasi et al., 2004) and dunites from Mauna Kea (Hawai'i) and La Réunion (Berger & Vannier, 1984; Vannier et al., 1982) also present evidence of ductile deformation such as subgrain boundaries and CPOs, but they are not well documented and quantified as in our study. Note that in Tommasi et al. (2004), the olivines display Mg-Fe gradients between cores (Fo# 89) and rims (Fo#82), thus permitting to evidence and constrain the extent of melt-rock interaction. Such investigation could not be achieved for Kaupulehu olivines as the Mg-Fe composition was completely re-equilibrated (no profile, Figure S6 and Table S5 in Supporting Information S1).

One can question how the internal misorientation and subgrain boundaries could have been preserved in Kaupulehu olivines at high temperature. Experiments constraining the kinetics of dislocation recovery (Farla et al., 2011) reported very slow recovery rates, implying that over 200 Ma are needed to fully annihilate both mobile and locked dislocations along subgrain boundaries in a 1-mm olivine. Microstructural indicators such as subgrain boundaries and CPOs can thus be preserved, while geochemical indicators such as Fo# and some minor element concentrations can be entirely reset.

At last, another hypothesis needs to be examined. Internal misorientation and apparent subgrain boundaries have been previously reported in magmatic euhedral olivines (Clague & Denlinger, 1994; Welsch et al., 2014; Wieser et al., 2020). These studies proposed that the microstructural features can be generated by ductile deformation in the crystal-mush pile after olivines settle and interconnect within a magmatic reservoir (Clague & Denlinger, 1994; Wieser et al., 2020). Grain interconnection in the pile would lead to a local high stress field, which could generate dislocation glides into the olivine lattice. Nevertheless, they reported very low internal misorientation across subgrain boundaries, up to 3° only (Wieser et al., 2020), with no CPO or SPO since they studied crushed tephra. Their interpretation is not compatible with the high values of internal misorientation (up to 18°) and strong CPOs in Kaupulehu olivine.

### 4.3. Synthesis

A petrogenetic scenario can be proposed based on our results. Orthopyroxene-rich mantle peridotites were first infiltrated by a reactive melt. Pyroxenes, especially orthopyroxenes, were dissolved at high temperatures (at least at 1,200°C), while olivines, previously deformed in the convecting mantle, were mostly preserved. Upon cooling, the intergranular clinopyroxenes and plagioclases crystalized, leading to the poikilitic texture. Concomitantly, the deformed olivines underwent a geochemical resetting by ionic diffusion, leading to homogenous compositions in major and most minor elements, that are similar to magmatic olivine (e.g., Fo# 84.2 on average) and almost complete dehydration. Geochemical resetting remained incomplete for some FRTEs whose compositions lie between those of mantle-derived olivines and OIB olivines (Figures 5 and 7f). Re-equilibration with the melt also induced partial dissolution of olivine edges as observed in olivines within the clinopyroxene oikocrysts (Figure S3 in Supporting Information S1, NBH19 zoom). Furthermore, this event was not long enough, or temperature was not high enough, to induce a complete recovery of the microstructures (subgrain boundaries walls, static recrystallization until equilibrium, fading of the CPOs) (Figure 3 and Figure S3 in Supporting Information S1). Finally, during further cooling and ascent toward the surface, volatiles exsolved leading to bubbles and residual SiO₂-rich melt solidified as glass films and small crystals along grain boundaries.

## 5. Conclusions

The six ultra-mafic nodule beds from the Kaupulehu lava flow of Hualālai volcano have antagonistic features. A magmatic origin for pyroxenes is supported by their poikilitic texture in NBH8, NBH16 and NBH19, and by the low olivine Fo# (80.8–87.9). A cumulative origin is further confirmed by the scarcity of orthopyroxenes in most samples, except for NBH16, by the co-existence of Fe-rich olivine and clinopyroxene, and by high TiO₂
concentration in spinel. The high equilibrium temperatures obtained for pyroxenes from NBH16 also advocate for a magmatic origin. For pyroxenes, a cumulative origin is logically proposed and follows the conclusion of previous studies (e.g., Chen et al., 1992; Gao et al., 2016).

However, based on our microstructural investigation, the petrogenesis of Kaupulehu olivines emerged as more complex. EBSD maps revealed olivines with well marked subgrain boundaries, internal misorientation up to 18°, and well developed CPO. The deformed olivines were homogeneously distributed in the samples. We also identified no SPO but a non-negligible recrystallized fraction of olivine (24%–42%). All evidence points toward olivines being inherited relics of a deformed mantle. Furthermore, trace and ultra-trace element compositions in olivine suggest at least a moderate melt-rock interaction leading to small enrichment in REE, although keeping similar patterns when compared to mantle compositions. Olivine compositions in Ni, Ca, and V likewise depict a composition in between OIB olivines and mantle olivines. The Kaupulehu olivine compositions and preservation of the microstructure can rationally be explained using laboratory-based ionic diffusion coefficients and dislocation recovery rates (e.g., Chakraborty, 2010; Farla et al., 2011). The different origins of pyroxenes and olivines are further supported by the estimated apparent equilibrium temperatures, which report a lower equilibrium temperature of olivine (Al-in-olivine: 1,000–1,100°C) than for pyroxenes (Ca in pyroxenes: 1,200°C). Following pioneer studies (e.g., Berger & Vannier, 1984; Tommasi et al., 2004; Vannier et al., 1982), we emphasize the crucial role of microstructural and EBSD investigation to infer the petrological origin of olivines and pyroxenes from cumulative ultra-mafic and mafic rocks.

## Data Availability Statement

The EBSD data (Dujardin et al., 2024a) and FTIR data (Dujardin et al., 2024b) supporting this study are openly available in the Zenodo repository.

## Acknowledgments

C. Nevado and D. Delmas are thanked for providing high-quality sections for FTIR and EBSD; J.L. Devidal for assistance with EPMA analyses; A. Tommasi and B. Ildefonse are deeply thanked for scientific discussions at an early stage of this project. The EBSD-SEM national facility at Geoscience Montpellier, the EPMA and SEM-FIB-STEM CarMa facility at Laboratoire Magmas et Volcans are all supported by the Institut National de Sciences de l'Univers (INSU), Centre National de la Recherche Scientifique (CNRS, France). We acknowledge funding from ANR H-Deep-ISO to S. Demouchy, from ANU-Research School of Earth Sciences to O. Alard, and from ARC-DP220100709 to O. Alard and S. Demouchy, from program TelluS of the Institut National des Sciences de l'Univers — CNRS, from PHC FASIC ECPD 2025 project N°53572VF, from the French government initiative "programme d'investissements d'avenir" managed by the Agence Nationale de la Recherche (ANR) under Grant ANR-20-SFRI-0003 project CAP GS, and from the University Clermont Auvergne via its Centre International de Recherche Catastrophes Naturelles et Développement Durable (CIR4) program (InMelt project) to S. Demouchy. This is CIR4-ClerVolc publication N°678.

## References

Arai, S. (1994). Characterization of spinel peridotites by olivine-spinel compositional relationship: Review and interpretation. *Chemical Geology*, **113**(3), 191–204. https://doi.org/10.1016/0009-2541(94)90066-3

Bachmann, F., Hielcher, R., & Schaeben, H. (2011). Grain detection from 2D and 3D EBSD data specification of the MTEX algorithm. *Ultramicroscopy* (Amsterdam), **111**(12), 1720–1733. https://doi.org/10.1016/j.ultramic.2011.08.002

Batanova, V. G., Thompson, J. M., Danyushevsky, L. V., Portnyagin, M. V., Garbe-Schönberg, D., Hauri, E., et al. (2019). New olivine reference material for in situ microanalysis. Geostand. *Geostandards and Geoanalytical Research*, **43**(3), 453–473. https://doi.org/10.1111/ggr.12266

Bell, D. R., Ihinger, P. D., & Rossman, G. R. (1995). Quantitative analysis of trace OH in garnet and pyroxenes. *American Mineralogist*, **80**(5–6), 465–474. https://doi.org/10.2138/am-1995-5-607

Bell, D. R., Rossman, G. R., Maldener, J., Endisch, D., & Rauch, F. (2003). Hydroxide in olivine: A quantitative determination of the absolute amount and calibration of the IR spectrum. *Journal of Geophysical Research*, **108**(B2). https://doi.org/10.1029/2001JB000679

Berger, E. T., & Vannier, M. (1984). Les dunites en enclaves dans les basaltes alcalins des îles océaniques: Approche pétrologique. *Bulletin de Mineralogie*, **107**(5), 649–663. https://doi.org/10.3406/bulmi.1984.7809

Blundy, J., & Wood, B. (2003). Partitioning of trace elements between crystals and melts. *Earth and Planetary Science Letters*, **210**(3–4), 383–397. https://doi.org/10.1016/S0012821X(03)00129-8

Bodinier, J.-L., & Godard, M. (2014). Orogenic, ophiolitic, and abyssal peridotites. In H. D. Holland & K. K. Turekian (Eds.), *Treatise on Geochemistry* (2nd ed.) (Vol. 3, pp. 103–167). Elsevier.

Bohrson, W. A., & Clague, D. A. (1988). Origin of ultramafic xenoliths containing exsolved pyroxenes from Hualalai volcano, Hawaii. *Contributions to Mineralogy and Petrology*, **100**(2), 139–155. https://doi.org/10.1007/BF00373581

Boudier, F., & Nicolas, A. (1995). Nature of the Moho transition zone in the Oman ophiolite. *Journal of Petrology*, **36**(3), 777–796. https://doi.org/10.1093/petrology/36.3.777

Bouvet de Maisonneuve, F., Costa, F., Huber, C., Vonlanthen, P., Bachmann, O., & Dungan, M. A. (2016). How do olivines record magmatic events? Insights from major and trace element zoning. *Contributions to Mineralogy and Petrology*, **171**(6), 56. https://doi.org/10.1007/s00410-016-1264-6

Brey, G. P., & Köhler, T. (1990). Geothermobarometry in four-phase lherzolite II. New thermobarometer, and practical assessment of existing thermobarometer. *Journal of Petrology*, **31**(6), 1353–1378. https://doi.org/10.1093/petrology/31.6.1353

Bunge, H.-J. (1982). *Texture analysis in materials science: Mathematical Methods* (p. 593). Butterworths.

Bussweiler, Y., Brey, G. P., Pearson, D. G., Stachel, T., Stern, R. A., Hardman, M. F., et al. (2017). The aluminum-in-olivine thermometer for mantle peridotites—experimental versus empirical calibration and potential applications. *Lithos*, **272–273**, 301–314. https://doi.org/10.1016/j.lithos.2016.12.015

Bussweiler, Y., Giuliani, A., Greig, A., Kjarsgaard, B. A., Petts, D., Jackson, S. E., et al. (2019). Trace element analysis of high-Mg olivine by LA-ICP-MS—characterization of natural olivine standards for matrix-matched calibration and application to mantle peridotites. *Chemical Geology*, **524**, 136–157. https://doi.org/10.1016/j.chemgeo.2019.06.019

Chakraborty, S. (2010). Diffusion coefficients in olivine, wadsleyite and ringwoodite. In *Reviews in Mineralogy and Geochemistry* (Eds.). Y. Zhang & D. Cherniak, American Mineralogical Society, Geochemical Society, (Vol. 72(1), 603–639). https://doi.org/10.2138/rmg.2010.72.13

Chen, C., Presnall, D. C., & Stern, R. J. (1992). Petrogenesis of ultramafic xenoliths from the 1800 Kaupulehu flow, Hualalai volcano, Hawaii. *Journal of Petrology*, **33**(1), 163–202. https://doi.org/10.1093/petrology/33.1.163

Chen, Y., & Zhang, Y. X. (2008). Olivine dissolution in basaltic melt. *Geochimica et Cosmochimica Acta*, **72**(19), 4756–4777. https://doi.org/10.1016/j.gca.2008.07.014
Cherniak, D. J. (2010). REE diffusion in olivine. *American Mineralogist*, 95(2-3), 362-368. https://doi.org/10.2138/am.2010.3345

Chin, E. J. (2018). Deep crustal cumulates reflect patterns of continental rift volcanism beneath Tanzania. *Contributions to Mineralogy and Petrology*, 173(10), 85. https://doi.org/10.1007/s00410-018-1512-z

Clague, D. A. (1982). Petrology of tholeiitic basalt dredged from Hualalai Volcano, Hawaii. *Eos, American Geophysical Union Transaction*, 63(45), 1138.

Clague, D. A. (1987). Hawaiian xenolith populations, magma supply rates, and development of magma chambers. *Bulletin of Volcanology*, 49(4), 577-587. https://doi.org/10.1007/bf01079963

Clague, D. A., & Denlinger, R. P. (1994). Role of olivine cumulates in destabilizing the flanks of Hawaiian volcanoes. *Bulletin of Volcanology*, 56(6-7), 425-434. https://doi.org/10.1007/BF00302824

Clague, D. A., Jackson, E. D., & Wright, T. L. (1980). Petrology of Hualalai volcano, Hawaii: Implication for mantle composition. *Bulletin of Volcanology*, 43(4), 641-656. https://doi.org/10.1007/BF02600363

Coogan, L. A. C., Hain, A., Stahk, S., & Chakraborty, S. (2005). Experimental determination of the diffusion coefficient for calcium in olivine between 900°C and 1,500°C. *Geochimica et Cosmochimica Acta*, 69(14), 3683-3694. https://doi.org/10.1016/j.gca.2005.03.002

Deer, W. A., Howie, R. A., & Zussman, J. (1997). *Rock forming minerals: Orthosilicates* (2nd ed.). (Vol. 1A). The Geological Society.

De Hoog, J. C. M., Gall, L., & Cornell, D. H. (2010). Trace-element geochemistry of mantle olivine and application to mantle petrogenesis and geothermobarometry. *Chemical Geology*, 270(1-4), 196-215. https://doi.org/10.1016/j.chemgeo.2009.11.017

Demouchy, S. (2010). Diffusion of hydrogen in olivine grain boundaries and implications for the survival of water-rich zones in the Earth's mantle. *Earth and Planetary Science Letters*, 295(1-2), 305-313. https://doi.org/10.1016/j.epsl.2010.04.019

Demouchy, S., & Alard, O. (2021). Hydrogen, trace, and ultra-trace element distribution in natural olivines. *Contributions to Mineralogy and Petrology*, 176(4), 26. https://doi.org/10.1007/s00410-021-01778-5

Demouchy, S., Barou, F., Ishikawa, A., Gardés, E., & Tommasi, A. (2024). Microstructures, water contents, and seismic properties of a tectonically exhumed sliver of oceanic lithosphere from the Moa Island, Timor-Tanimbar outer-arc, eastern Indonesia. *Tectonophysics*, 887, 230443. https://doi.org/10.1016/j.tecto.2024.230443

Demouchy, S., & Bolfan-Casanova, N. (2016). Distribution and transport of hydrogen in the lithospheric mantle: A review. *Lithos*, 240, 402-425. https://doi.org/10.1016/j.lithos.2015.11.012

Demouchy, S., Ishikawa, A., Tommasi, A., Alard, O., & Keshav, S. (2015). Characterization of hydration in the mantle lithosphere: Peridotite xenoliths from the Ontong Java Plateau as an exemple. *Lithos*, 212, 189-201. https://doi.org/10.1016/j.lithos.2014.11.005

Demouchy, S., & Mackwell, S. (2006). Mechanisms of hydrogen incorporation and diffusion in iron-bearing olivine. *Physics and Chemistry of Minerals*, 33(5), 347-355. https://doi.org/10.1007/s00269-006-0081-2

Demouchy, S., Shcheka, S., Denis, C. M. M., & Thoraval, C. (2017). Subsolidus hydrogen partitioning between nominally anhydrous minerals in garnet-bearing peridotite. *American Mineralogist*, 102(9), 1822-1831. https://doi.org/10.2138/am-2017-6089

Demouchy, S., & Tommasi, A. (2021). From dry damp and stiff mantle lithosphere by reactive melt percolation atop the Hawaii plume. *Earth and Planetary Science Letters*, 574, 117159. https://doi.org/10.1016/j.epsl.2021.117159

Demouchy, S., Tommasi, A., Ionov, D., Higgie, K., & Carlson, R. W. (2019). Microstructures, water contents, and seismic properties of the mantle lithosphere beneath the northern limit of the Hangay Dome, Mongolia. *Geochemistry, Geophysics, Geosystems*, 20(1), 183-207. https://doi.org/10.1029/2018GC007931

Demouchy, S., Wang, Q., & Tommasi, A. (2023). Deforming the upper Mantle—olivine mechanical properties and anisotropy. *Elements*, 19(3), 151-157. https://doi.org/10.2138/gselements.19.3.151

Denis, C. M. M., Alard, O., & Demouchy, S. (2015). Water content and hydrogen behaviour during metasomatism in the uppermost mantle beneath Ray Pic volcano (Massif Central, France). *Lithos*, 236-237, 256-274. https://doi.org/10.1016/j.lithos.2015.08.013

Dixon, J. E., Leist, L., Langmuir, C., & Schilling, J.-G. (2002). Recycled dehydrated lithosphere observed in plume-influenced mid-ocean-ridge basalt. *Nature*, 420(6914), 385-389. https://doi.org/10.1038/nature01215

Dohmen, R., Becker, H. W., & Chakraborty, S. (2007). Fe-Mg diffusion in olivine I. Experimental determination between 700 and 1,200°C as a function of composition, crystal orientation and oxygen fugacity. *Physics and Chemistry of Minerals*, 34(6), 389-407. https://doi.org/10.1007/s00269-007-0157-7

Drouin, M., Godard, M., Ildefonse, B., Bruguer, O., & Garrido, C. J. (2009). Geochemical and petrographic evidence for magmatic impregnation in the ocean lithosphere at Atlantis Massif, Mid-Atlantic Ridge (IODP Hole U1309D, 30°N). *Chemical Geology*, 264(1-4), 71-88. https://doi.org/10.1016/j.chemgeo.2009.02.013

Drouin, M., Ildefonse, B., & Godard, M. (2010). A microstructural imprint of melt impregnation in slow spreading lithosphere: Olivine-rich troctolites from the Atlantis Massif, Mid-Atlantic Ridge, 30°N, IODP Hole U1309D. *Geochemistry, Geophysics, Geosystems*, 11(6). https://doi.org/10.1029/2009GC002995

Dujardin, A. D., Demouchy, S., Alard, O., Gardés, E., Laubier, M., & Barou, F. (2024a). Kaupulehu peridotites—EBSD-CTF files [Dataset]. *Zenodo*. https://doi.org/10.5281/zenodo.14418048

Dujardin, A. D., Demouchy, S., Alard, O., Gardés, E., Laubier, M., & Barou, F. (2024b). Kaupulehu peridotites—Mid FTIR—Unpolarized average spectra files [Dataset]. *Zenodo*. https://doi.org/10.5281/zenodo.14418271

Farla, R. J. M., Kokkonen, H., Fitz Gerald, J. D., Barnhoorn, A., Faul, U. H., & Jackson, I. (2011). Dislocation recovery in fine-grained polycrystalline olivine. *Physics and Chemistry of Minerals*, 38(5), 363-377. https://doi.org/10.1007/s00269-010-0410-3

Férot, A., & Bolfan-Casanova, N. (2012). Water storage capacity in olivine and pyroxene to 14 GPa: Implications for the water content of the Earth's upper mantle and nature of seismic discontinuities. *Earth and Planetary Science Letters*, 349-350, 218-230. https://doi.org/10.1016/j.epsl.2012.06.022

Fodor, R. V., & Galar, P. (1997). A view into the subsurface of Mauna Kea volcano, Hawaii: Crystallization processes interpreted through the petrology and petrography of gabbroic and ultramafic xenoliths. *Journal of Petrology*, 38(5), 581-624. https://doi.org/10.1093/petroj/38.5.581

Foley, S. F., Prelevic, D., Rehfeldt, T., & Jacob, D. E. (2013). Minor and trace elements in olivines as probes into early igneous and mantle melting processes. *Earth and Planetary Science Letters*, 363, 181-191. https://doi.org/10.1016/j.epsl.2012.11.025

Gao, R., Lassiter, J. C., Barnes, J. D., Clague, D. A., & Bohrson, W. A. (2016). Geochemical investigation of Gabbroic Xenoliths from Hualalai Volcano: Implications for lower oceanic crust accretion and Hualalai Volcano magma storage system. *Earth and Planetary Science Letters*, 442, 162-172. https://doi.org/10.1016/j.epsl.2016.02.043

Gao, R., Lassiter, J. C., Clague, D. A., & Bohrson, W. A. (2022). Evolution of Hawaiian volcano magmatic plumbing system and implications for melt/edifice and melt/lithosphere interaction: Constraints from Hualalai xenoliths. *Journal of Petrology*, 63(9), 1-24. https://doi.org/10.1093/jpetrology/egac091

Griffin, W. L., Powell, W., Pearson, N. J., & O'Reilly, S. Y. (2008). Glitter: Data reduction software for laser ablation ICP-MS. Laser ablation-ICP-MS in the earth Sciences. *Mineralogical Association of Canada Short Course Series*, 40(Appendix 2), 204-207.
Guillong, M., & Heinrich, C. A. (2007). Sensitivity enhancement in laser ablation ICP-MS using small amounts of hydrogen in the carrier gas. *Journal of Analytical Atomic Spectrometry*, 22(12), 1488–1494. https://doi.org/10.1039/B709489B

Hirschmann, M. M., Tenner, T., Aubaud, C., & Wither, A. C. (2009). Deshydration melting of nominally anhydrous mantle: The primacy of partitioning. *Physics of the Earth and Planetary Interiors*, 176(1–2), 54–68. https://doi.org/10.1016/j.pepi.2009.04.001

Holtzman, B. K., Kohlstedt, D. L., Zimmerman, M. E., Heidelbach, F., Hiraga, T., & Hustoft, J. (2003). Melt segregation and strain partitioning: Implications for the seismic anisotropy and mantle flow. *Science*, 301(5637), 1227–1230. https://doi.org/10.1126/science.1087132

Ishikawa, A., Maruyama, S., & Komiya, T. (2004). Layered lithospheric mantle beneath the ontong java plateau: Implications from xenoliths in alnoite, Malai'a, Salomon islands. *Journal of Petrology*, 45(10), 2011–2044. https://doi.org/10.1093/petrology/egh046

Jackson, E. D., Clague, D. A., Engleman, E., Friesen, W. F., & Norton, D. (1981). *Xenoliths in the Alkalic basalt flows from Hualalai volcano, Hawaii (Report 81-1031)*. U. S. Geological Survey.

Jochum, K. P., Nohl, U., Herwig, K., Lammel, E., Stoll, B., & Hofmann, A. W. (2005). GeoReM: A new geochemical database for reference materials and isotopic standards. *Geostandards and Geoanalytical Research*, 29(3), 333–338. https://doi.org/10.1111/j.1751-908X.2005.tb00904.x

Jung, H., & Karato, S. (2001). Water-induced fabric transition in olivine. *Science*, 293(5534), 1460–1463. https://doi.org/10.1126/science.1062235

Kirby, S. H., & Green, H. W. (1980). Dunite xenoliths from Hualalai volcano: Evidence for mantle diapiric flow beneath the island of Hawaii. *American Journal of Science*, 280, 550–575.

Lambart, S., Hamilton, S., & Lang, O. I. (2022). Compositional variability of San Carlos olivine. *Chemical Geology*, 605, 120968. https://doi.org/10.1016/j.chemgeo.2022.120968

Le Maitre, R. W., Streckeisen, A., Zanettin, B., Le Bas, M. J., Bonin, B., Bateman, P., & Lameyre, J. (2002). Igneous rocks: A classification and glossary of terms (2nd ed.). Cambridge University Press. https://doi.org/10.1017/CBO9780511535581

Lin, K.-Y., Warren, J. M., & Davis, F. A. (2023). Trace elements in abyssal peridotite olivine record melting, thermal evolution, and melt refertilization in the oceanic upper mantle. *Contributions to Mineralogy and Petrology*, 178(10), 66. https://doi.org/10.1007/s00410-023-02044-6

Lopez-Sanchez, M. A., Tommasi, A., Ben Ismail, W., & Barou, F. (2021). Dynamic recrystallization by subgrain rotation in olivine revealed by electron backscatter diffraction. *Tectonophysics*, 815, 228916. https://doi.org/10.1016/j.tecto.2021.228916

Mainprice, D., Bachmann, F., Hielscher, R., & Schaeben, H. (2014). Descriptive tools for the analysis of texture projects with large datasets using MTEX: *Strength, symmetry, and component*. (Vol. 409, pp. 251–271). Geological Society, London, Special Publication. https://doi.org/10.1144/SP409.8

McDonough, W., & Sun, S.-S. (1995). The composition of the Earth. *Chemical Geology*, 120(3–4), 223–253. https://doi.org/10.1016/0009-2541(94)00140-4

Mercier, J.-C., & Nicolas, A. (1975). Textures and fabrics of upper-mantle peridotites as illustrated by xenoliths from basalts. *Journal of Petrology*, 16(1), 454–487. https://doi.org/10.1093/petrology/16.1.454

Moore, R. B., Clague, D. A., Rubin, M., & Bohrson, W. A. (1987). Hualalai volcano: A preliminary summary of geologic, petrologic and geophysical data. In R. W. Decker, T. L. Wright, & P. H. Stauffer (Eds.), *Volcanism in Hawaii* (Vol. 1). U.S. Geological Survey professional. https://doi.org/10.3133/pp1350

Mourey, A. J., Shea, T., Lynn, K. J., Lerner, A. H., Lambart, S., Costa, F., et al. (2022). Trace elements in olivine fingerprint the source of 2018 magmas as shed light on explosive-effusive eruption cycles at Kilauea Volcano. *Earth and Planetary Science Letters*, 595, 117769. https://doi.org/10.1016/j.epsl.2022.117769

Mutch, E. J. F., Maclennan, J., Schortle, O., Edmonds, M., & Rudge, J. F. (2019). Rapid transcrustal magma movement under Iceland. *Nature Geoscience*, 12, 569–574. https://doi.org/10.1038/s41561-019-0376-9

Neumann, E. R., Wulff-Pedersen, E., Simonsen, S. L., Pearson, N. J., Marti, J., & Mitjavila, J. (1999). Evidence for fractional crystallization of periodically refilled magma chambers in Tenerife, Canary Islands. *Journal of Petrology*, 40(7), 1089–1123. https://doi.org/10.1093/petroj/40.7.1089

O'Reilly, S. Y., Chen, D., Griffin, W. L., & Ryan, C. G. (1997). Minor elements in olivine from spinel lherzolite xenoliths: Implications for thermobarometry. *Mineralogical Magazine*, 61(405), 257–269. https://doi.org/10.1180/minmag.1997.061.405.09

O'Reilly, S. Y., & Griffin, W. L. (2013). Mantle metasomatism. In D. E. Harlov & H. Austrheim (Eds.), *Metasomatism and the chemical transformation of rock: The role of fluids in terrestrial and extraterrestrial processes* (pp. 471–533). Nature. https://doi.org/10.1007/978-3-642-28394-9_12

Paterson, M. S. (1982). The determination of hydroxyl by infrared absorption in quartz, silicate glasses and similar materials. *Bulletin de Mineralogie*, 105(1), 20–29. https://doi.org/10.3406/bulmi.1982.7582

Peslier, A. H., & Bizimis, M. (2015). Water in Hawaiian peridotite minerals: A case for a dry metasomatized oceanic mantle lithosphere. *Geochemistry, Geophysics, Geosystems*, 16(4), 1211–1232. https://doi.org/10.1002/2015GC005780

Petry, C., Chakraborty, S., & Palme, H. (2004). Experimental determination of Ni diffusion coefficient in olivine and their dependence on temperature, composition, oxygen fugacity, and crystallographic orientation. *Geochimica et Cosmochimica Acta*, 68(20), 4179–4188. https://doi.org/10.1016/j.gca.2004.02.024

Richter, D. H., & Murata, K. J. (1961). *Xenolithic nodules in the 1800-1801 Kaupulehu flow of Hualalai volcano, Hawaii* (Vol. 424-B, pp. B215-B217). U.S. Geological Survey Professional Paper.

Roedder, E. (1965). Liquid CO₂ inclusions in olivine-bearing nodules and phenocrysts from basalts. *American Mineralogist*, 50(10), 1746–1782.

Satsukawa, T., Gadard, M., Demouchy, S., Michibayashi, K., & Ildefonse, B. (2017). Chemical interactions in the subduction factory: New insight from an in situ trace element and hydrogen study of the Ichinomegata and Oki-Dogo mantle xenoliths (Japan). *Geochimica et Cosmochimica Acta*, 208, 234–267. https://doi.org/10.1016/j.gca.2017.03.042

Sen, G., Frey, F. A., Shimizu, N., & Leeman, W. P. (1993). Evolution of the lithosphere beneath Oahu, Hawaii: Rare Earth element abundance in mantle xenoliths. *Earth and Planetary Science Letters*, 119(1–2), 53–69. https://doi.org/10.1016/0012-821X(93)90006-U

Sjoblom, M. P., Dorais, M. J., Christiansen, E. H., & Fodor, R. V. (2023). Mafic to ultramafic xenoliths from Mauna Kea, Hawaii: Clue to magma evolution from trace element composition of clinopyroxene. *Contributions to Mineralogy and Petrology*, 178(3), 20. https://doi.org/10.1007/s00410-023-02001-3

Skogby, H. (2006). Water in natural mantle minerals I: Pyroxenes. *Mineralogy and Geochemistry*, 62(1), 155–167. https://doi.org/10.2138/rmg.2006.62.7

Sobolev, A. V., Hofmann, A. W., Kuzmin, D. V., Yaxley, G. M., Arndt, N. T., Chung, S. L., et al. (2007). The amount of recycled crust in source of mantle-derived melts. *Science*, 316(5823), 412–417. https://doi.org/10.1126/science.1138113
Spandler, C., & O'Neill, H. S. C. (2010). Diffusion and partition coefficients of minor and trace elements in San Carlos olivine at 1,300°C with some geochemical implications. *Contributions to Mineralogy and Petrology*, **159**(6), 791–818. https://doi.org/10.1007/s00410-009-0456-8

Spandler, C., O'Neill, H. S. C., & Kamenetsky, V. S. (2007). Survival times of anomalous melt inclusions from element diffusion in olivine and chromite. *Nature*, **447**(7142), 303–306. https://doi.org/10.1038/nature05759

Stearns, H. T., & Macdonald, G. A. (1946). *Geology and ground-water resources of the island of Hawaii. Hawaii division of hydrography* (*bulletin* 9 (p. 363). U.S. Geological Survey.

Thoraval, C., Demouchy, S., & Padrón-Navarta, J. A. (2019). Relative diffusivities of hydrous defectes from a partially dehydrated natural olivine. *Physics and Chemistry of Minerals*, **46**(1), 1–13. https://doi.org/10.1007/s00269-018-0982-x

Tommasi, A., Godard, M., Coromina, G., Dautria, J. M., & Barsczus, H. (2004). Seismic anisotropy and compositionally induced velocity anomalies in the lithosphere above mantle plume: A petrological and microstructural study of mantle xenoliths from French Polynesia. *Earth and Planetary Science Letters*, **227**(3–4), 539–556. https://doi.org/10.1016/j.epsl.2004.09.019

Tommasi, A., Mainprice, D., Canova, G., & Chastel, Y. (2000). Viscoplastic self-consistent and equilibrium-based modeling of olivine lattice preferred orientations: Implication for the upper mantle seismic anisotropy. *Journal of Geophysical Research*, **105**(B4), 7893–7908. https://doi.org/10.1029/1999JB900411

Tommasi, A., Mameri, L., & Godard, M. (2020). Textural and compositional changes in the lithospheric mantle atop the Hawaiian plume: Consequences for seismic properties. *Geochemistry, Geophysics, Geosystems*, **21**(8). https://doi.org/10.1029/2020GC009138

Vannier, M., Besson, M., & Berger, E. T. (1982). Remarque sur la genèse et la cristallisation des magmas basaltique. *Bulletin de Mineralogie, 105*(4), 376–385. https://doi.org/10.3406/bulmi.1982.7632

Vernon, R. H. (2000). *Review of microstructural evidence of magmatic and solid-state flow*. Electronic Geosciences 5.

Veter, M., Foley, S. F., Mertz-Kraus, R., & Groschopf, N. (2017). Trace elements in olivine of ultramafic lamprophyres controlles by phlogopite-rich mineral assemblages in the mantle source. *Lithos*, **292**, 81–95.

Wang, C., Liang, Y., Xu, W., & Dygert, N. (2013). Effect of melt composition on basalt and peridotite interaction: Laboratory dissolution experiment with applications to mineral compositional variations in mantle xenoliths from the north China craton. *Contributions to Mineralogy and Petrology*, **166**(5), 1469–1488. https://doi.org/10.1007/s00410-013-0938-6

Welsch, B., Hammer, J., & Hellebrand, E. (2014). Phosphorus zoning reveals dendritic architecture of olivine. *Geology*, **42**(10), 867–870. https://doi.org/10.1130/G35691.1

White, R. W. (1966). Ultramafic inclusions in basaltic rocks from Hawaii. *Contributions to Mineralogy and Petrology*, **12**(3), 245–314. https://doi.org/10.1007/bf00518082

Wieser, P. E., Edmonds, M., Maclennan, J., & Wheeler, J. (2020). Microstructural constraints on magmatic mushes under Kilauea volcano, Hawaii. *Nature Communications*, **11**(1), 14. https://doi.org/10.1038/s41467-019-13635-y

Withers, A. C., Bureau, H., Raepsaet, C., & Hirschmann, M. M. (2012). Calibration of infrared spectroscopy by elastic recoil detection analysis of H in synthetic olivine. *Chemical Geology*, **334**, 92–98. https://doi.org/10.1016/j.chemgeo.2012.10.002

Ziberna, L. (2021). Geothermobarometry of mafic and ultramafic xenoliths: Examples from hualalai and Mauna Kea volcanoes, Hawaii. In M. Masotta, C. Beier, & S. Mollo (Eds.), *Crustal magmatic system evolution: Anatomy, architecture, and physico-chemical processes* (Vol. 264, pp. 1–18). Geophysical Monograph Series. https://doi.org/10.1002/9781119564485.ch1