Figure 1. Geological maps of the Isua supracrustal belt,
southwestern Greenland and north-central portion of the East Pilbara
Terrane, northwestern Australia. a: simplified geology of the
Isua supracrustal belt and adjacent areas [modified from Nutman et al.
(2002)]. Locations of meta-peridotite enclaves and lenses A and B are
presented. b: simplified geology of the north edge of the Mount
Edgar Complex [modified from Van Kranendonk et al. (2007)] showing
major km-scale ultramafic intrusive bodies: the Gap Intrusion, the Nob
Webb Intrusion, and the Strutton Intrusion. c: location of the
Isua supracrustal belt in southwestern Greenland. d: location
of the East Pilbara Terrane in northwestern Australia. Yellow circles:
locations for new samples; white circles, locations for compiled samples
from Szilas et al. (2015), Van de Löcht et al. (2018), Friend et al.
(2002), Friend and Nutman (2011), McIntyre et al. (2019), Dymek et al.
(1988). and the Geological Survey of Western Australia 2013 database.
Geological background and proposed tectonic models
The Isua supracrustal belt
The ~35 km-long, ~1 to 4 km-wide Isua
supracrustal belt of southwestern Greenland is Earth’s largest
recognized Eoarchean terrane (Fig. 1a ). The belt’s protoliths
were formed during ~3.8 Ga to ~3.7 Ga
and experienced extensive shearing, thinning, and folding (e.g., Nutman
et al., 2020; Webb et al., 2020). Regional deformation of the Isua
supracrustal belt is associated with amphibolite facies assemblages that
have been interpreted to be Eoarchean (e.g., Nutman et al., 2020; Webb
et al., 2020; Ramirez-Salazar et al., 2021; Zuo et al., 2021) and/or
Neoarchean in age (e.g., Chadwick, 1990; Nutman, 1986; Nutman et al.,
2015). Meta-tonalites of similar ages to the ~3.8 to 3.7
Ga supracrustal rocks are in contact with the Eoarchean supracrustal
belt to the north and south (Crowley et al., 2002; Crowley, 2003). The
interior of the belt exposes metamorphosed basalts chert, banded iron
formation, and voluminously minor metamorphosed ultramafic igneous
rocks, felsic volcanic rocks, and detrital sedimentary rocks (e.g.,
Komiya et al., 1999; Nutman et al., 2002; Nutman and Friend, 2009)
Ultramafic rocks in the Isua supracrustal belt occur as
~1- to ~100-m-scale lenticular bodies
associated with mafic pillow lavas (e.g., Dymek et al., 1988b; Szilas et
al., 2015) or as enclaves in both north and south meta-tonalite bodies
(e.g., Friend et al., 2002; Nutman and Friend, 2009). These ultramafic
rocks appear to have experienced various degrees of alteration including
carbonitization and serpentinization (e.g., Dymek et al., 1988b; Friend
et al., 2002; Szilas et al., 2015). Two
~104 m2meta-peridotite lenses (a southern lens A and a northern lens B) located
~1.5-km apart along the eastern edge of the western Isua
supracrustal belt and some ultramafic enclaves (as large as
~104 m2) in
meta-tonalite located ~15 km south of the belt
(Fig. 1a ) contain dunites and/or harzburgites with relatively
weak carbonitization and serpentinization. Igneous, metamorphic and
deformation features of these dunites and harzburgites have been
explored to constrain the Eoarchean tectonic evolution of the Isua
supracrustal belt (e.g., Kaczmarek et al., 2016; Nutman et al., 2020;
Van de Löcht et al., 2018). These features include: (1) primary rock
textures and deformation overprints, such as polygonal textures and
B-type olivine deformation fabrics observed in dunites from the
meta-peridotite lenses A and B in the Isua supracrustal belt (Kaczmarek
et al., 2016; Nutman et al., 1996); (2) a mineral assemblage of olivine
+ serpentine ± pyroxene ± Ti-humite ± carbonate ± spinel ± ilmenite ±
magnesite for dunites from lenses A and B (e.g., Guotana et al., 2021;
Nutman et al., 2020; Szilas et al., 2015 and a mineral assemblage of
olivine + serpentine + pyroxene + spinel ± hornblende for
meta-peridotites from the ultramafic enclaves (Van de Löcht et al.,
2018, 2020); (3) primitive mantle-normalized rare earth element patterns
(REE) that are sub-parallel to those of nearby basalts (e.g., Szilas et
al., 2015; Van de Löcht et al., 2020) or komatiite (Dymek et al.,
1988b); and (4) various highly siderophile element (HSE) patterns,
including relatively a high primitive mantle-normalized (PM-normalized)
Os, Ir and Ru versus to Pt and Pd pattern preserved in ultramafic
enclaves in the south meta-tonalite (Van de Löcht et al., 2018), and
similar or opposite patterns preserved in the two meta-peridotite lenses
of the Isua supracrustal belt (Szilas et al., 2015).
The Isua supracrustal belt has been mostly interpreted to record
~3.8 to 3.6 Ga plate tectonic processes including
subduction and subsequent extension (e.g., Arai et al., 2015; Komiya et
al., 1999; Nutman et al., 2020; Nutman et al., 2013b; Nutman and Friend,
2009). There are two main competing plate tectonic models for the
development of the belt that predict opposite subduction vergences
during the Eoarchean. One model that involves southward subduction
describes the belt as an Eoarchean accretionary prism, and has no
specific prediction for the generation and emplacement of the ultramafic
rocks (e.g., Arai et al., 2015; Komiya et al., 1999). In an alternative
model, the Isua supracrustal belt initially formed via intra-oceanic arc
magmatism during northward subduction and the subsequent collision of
multiple arc terranes at ~3.7 Ga (e.g., Nutman et al.,
2020). In this model, both the tonalites and supracrustal materials were
mostly generated by partial melting of materials from the mantle wedge,
subducting slab, or lowermost crust (e.g., Nutman et al., 2013a; Nutman
et al., 2020). The only exception is the presence of the dunite- and
harzburgite-hosting ultramafic enclaves and two lenses described above,
which have been proposed to represent relict melt-depleted mantle rocks
thrust atop crustal rocks in a subduction setting (see Figure 8 of
Nutman et al., 2013a). As such, (1) the geochemical associations with
local basalts are interpreted to reflect melt-rock reactions between
basaltic melts and depleted mantle residues (Friend and Nutman, 2011;
Van de Löcht et al., 2020); (2) a specific HSE pattern (i.e., relative
depletion of Pt, Pd, and Re versus Os, Ir, and Ru) found in ultramafic
enclaves enveloped by meta-tonalites located south of the Isua
supracrustal belt (Fig. 1a ) is thought to reflect fractionation
during melt depletion in the mantle (Van de Löcht et al., 2018); (3)
polygonal rock textures are interpreted to record equilibration under
mantle conditions (e.g., Nutman et al., 1996), whereas the B-type
olivine fabrics are claimed to exclusively indicate deformation in
hydrous mantle wedge environments (Kaczmarek et al., 2016, and
references therein); (4) the occurrence of an olivine + antigorite ±
Ti-humite mineral assemblage in some dunitic Isua ultramafic rocks is
interpreted as evidence of low-temperature, ultrahigh-pressure (UHP)
metamorphism (<500 °C , >2.6
GPa) that may be only compatible with a subduction setting (Nutman et
al., 2020); (5) the oxygen isotope signatures of some dunitic Isua
ultramafic rocks from lens A or lens B are considered to be indicative
of metasomatism by mantle-derived fluids, or of metamorphic growth of
olivine during the interpreted UHP metamorphism, respectively (Nutman et
al., 2021); and (6) clinopyroxene inclusions in olivine of Isua
ultramafic rocks from lens A are interpreted to represent
re-equilibration between ascending melts and melt-depleted mantle
peridotites (Nutman et al., 2021).
Recently, a heat-pipe model (i.e., a subcategory of hot stagnant-lid
tectonics) was proposed as an alternative to plate tectonics for the
formation and deformation of the Isua supracrustal belt (Webb et al.,
2020). Like other hot stagnant-lid tectonic models (e.g., Collins et
al., 1998; Johnson et al., 2014), heat-pipe tectonics is dominated by
(sub-)vertical transportation of materials, but the main driving force
of this transportation is volcanic advection rather than gravitational
instability (Moore and Webb, 2013; O’Reilly and Davies, 1981).
Voluminous mafic volcanism causes extensive volcanic resurfacing as well
as burial and downwards advection of cold surface materials. At great
depths, portions of buried hydrated mafic crust are partially melted,
forming tonalitic melts. Crustal deformation of a heat-pipe lithosphere
is predicted to happen via (1) radial shortening due to subsidence of
crustal materials in Earth’s quasi-spherical geometry (Bland and
McKinnon, 2016; Webb et al., 2020); or (2) contraction during a
plate-breaking and subduction event as or soon after the heat-pipe
cooling ceases (Beall et al., 2018; Webb et al., 2020). Alternatively,
deformation of a fragment of a heat-pipe lithosphere may be possible at
any time when involved in a younger deformation zone of any tectonic
setting. As to the formation of ultramafic rocks, this model does not
involve the thrusting of mantle rocks atop crustal rocks, as subduction
and associated mantle wedge settings do not occur during heat-pipe
cooling. In this model, the Eoarchean Isua supracrustal belt and
adjacent meta-tonalites were initially formed from ~3.8
to ~3.7 Ga via heat-pipe volcanism and lower crust
partial melting, and all ultramafic rocks would be crustal cumulates or
ultramafic lavas. The features of Isua ultramafic rocks which had
previously been interpreted in a plate tectonic context, as enumerated
in the prior paragraph, can be alternatively interpreted as follows: (1)
the geochemical relationships between Isua ultramafic rocks and nearby
basalts reflect contamination of ultramafic lava or cumulate mush by
co-existing fluids/melts; (2) the observed fractionated HSE patterns
were produced by partitioning during fractional crystallization; (3) the
rock textures of Isua ultramafic rocks were produced by crystallization
of melts and/or subsequent deformation/mineral re-equilibration under
crustal conditions; and (4) the mineral assemblages of these ultramafic
rocks, (5) fluid metasomatism, and (6) clinopyroxene inclusions in
olivine all formed under crustal conditions.
The East Pilbara Terrane
The ~40,000 km2 East Pilbara Terrane
of northwestern Australia is the largest and best-preserved Paleoarchean
terrane on Earth (Fig. 1b ). There, eleven granitoid bodies
(mostly meta-tonalites, with minor granites) are surrounded by broadly
coeval supracrustal belts. These supracrustal belts are dominantly
comprised of metamorphosed mafic to felsic volcanic rocks, with some
chemical and clastic sedimentary rocks, and ultramafic layered rocks and
intrusions (e.g., Van Kranendonk et al., 2007; Hickman, 2021). Rock
formation, deformation, and metamorphism (largely greenschist facies) in
the East Pilbara Terrane are thought to mostly occur from
~3.5 to 3.2 Ga, such that by the end of the
Paleoarchean, the supracrustal belts had been deformed into synforms and
the granitoids had become domes (Collins et al., 1998; Van Kranendonk et
al., 2007). This regional “dome-and-keel” geometry is a key element
for tectonic interpretations of the East Pilbara Terrane.
Ultramafic rocks of the East Pilbara Terrane occur as layers or pods
interleaved with supracrustal rocks or as km-scale igneous bodies
intruding supracrustal sequences (e.g., Smithies et al., 2007).
Ultramafic layers and pods found in the supracrustal sequences commonly
have thicknesses of ~1 to 5 meters and, preserve
spinifex textures in some locations. These rocks have been interpreted
to have been crystallized from komatiitic or basaltic lava flows (e.g.,
Smithies et al., 2007; Van Kranendonk et al., 2007). In this study, we
focus on the km-scale intrusions. In the East Pilbara Terrane,
ultramafic rocks are exposed as three >10-km-long and
>100-m-thick ultramafic intrusive bodies (Fig.
1b ), which include the Gap Intrusion, the Strutton Intrusion, and the
Nob Well Intrusion. These ultramafic bodies intrude
~3.53 to 3.43 Ga supracrustal sequences and are intruded
themselves by ~3.31 Ga granodiorites (Fig. 1b )
(Williams, 1999). Existing knowledge of these ultramafic rocks is mostly
limited to map relationships, petrological descriptions and geochemical
data published by the Geological Survey of Western Australia (e.g.,
Williams, 1999). In general, these ultramafic intrusions are comprised
of variably metamorphosed peridotite (including dunite), pyroxenite, and
gabbro (Geological Survey of Western Australia 2013 database).
It is now broadly accepted that the East Pilbara Terrane represents a
Paleoarchean terrane formed via regional hot stagnant-lid tectonics that
featured vigorous (ultra)mafic and felsic volcanism (e.g., Collins et
al., 1998; Johnson et al., 2017; François et al., 2014; Moore and Webb,
2013; Van Kranendonk et al., 2007; Van Kranendonk, 2010; Wiemer et al.,
2018). One subcategory of this tectonic regime is the partial convective
overturn cooling model, which was initially proposed based on the
geology of the East Pilbara Terrane (Collins et al., 1998). This model
predicts that the East Pilbara Terrane experienced episodic supracrustal
volcanism and tonalite formation followed by quiescence during
~10 to ~100 million years cycles of
mantle plume activities. (Ultra)mafic magmatism associated with mantle
plumes can produce km-scale ultramafic intrusions with or without
fractional crystallization (e.g., Smithies, 2007). The partial
convective overturn cooling model involves gravitational instability
between the relatively hot, buoyant tonalite bodies and colder, denser
supracrustal materials. Such instability could lead to diapiric rise of
tonalites, with supracrustal rocks deformed into synclines that wrap
around the tonalite domes, creating the observed “dome-and-keel”
geometry. No subduction activity and associated mantle-derived
ultramafic rocks are predicted at the crustal levels of a partial
convective overturn lithosphere (e.g., Collins et al., 1998).
Methods:
Three ultramafic samples (AL52614-4A, AW52614-4A, and AW52614-6)
collected from the Gap Intrusion of the East Pilbara Terrane and six
samples (AW17724-1, AW17724-2C, AW17724-4, AW17725-2B, AW17725-4 and
AW17806-1) collected from the Isua supracrustal belt were analyzed in
this study (Fig. 1 ). Isua samples AW17724-2C, AW17724-4
(northern lens B) and AW17725-4 (southern lens A) were collected from
the two meta-peridotite lenses which have been interpreted previously as
tectonic mantle slices (e.g., Friend and Nutman, 2011; Nutman et al.,
2020)]. Isua sample AW17724-1 was collected from the serpentinite
layer enveloping the meta-peridotite lens B. Isua sample AW17725-2B was
collected from an ultramafic outcrop near the northern meta-tonalite,
~300 meters east of the lens B. Isua sample AW17806-1
was collected from an outcrop located at the eastern supracrustal belt
near the northern meta-tonalite body (Fig. 1a, Table 1 ). To
test models of their petrogenesis, we compiled literature data and
inspected our samples using thin section petrography and acquisition of
(1) whole-rock major/trace element data (Table S1); (2) spinel
geochemistry (Table S2); and (3) HSE abundances (Table S3). Compiled
Isua and Pilbara data include results of previous studies focused on
ultramafic rocks located adjacent to our sample locations. These
outcrops specifically include (1) ultramafic rocks collected across the
Isua supracrustal belt (including the meta-peridotite lenses) studied by
Szilas et al. (2015) and Friend and Nutman (2011) (Fig. 1a );
(2) ultramafic rocks from the enclaves within the meta-tonalite located
south of the Isua supracrustal belt (Van de Löcht et al., 2018); and (3)
ultramafic rocks from the Nob Well Intrusion of the East Pilbara Terrane
(Geological Survey of Australia 2013 database; Fig. 1b ). Data
from other ultramafic rocks that have been variably interpreted as
cumulates or mantle peridotites (see Figures 3–8 captions for
references) are compiled for comparison with the ultramafic lithologies
of this study. These rocks were collected from variably deformed and
altered Archean ultramafic complexes (e.g., McIntyre et al., 2019),
massive layered intrusions (e.g., Coggon et al., 2015), collisional
massifs (e.g., Wang et al., 2013), volcanic xenoliths (e.g., Ionov,
2010) or mantle rocks extracted from ocean drilling (e.g., Parkinson and
Pearce, 2008).
Analytical details
The whole-rock major element concentrations of Pilbara ultramafic
samples were analyzed in the Peter Hooper GeoAnalytical Laboratory at
Washington State University. Major elements (e.g. MgO, FeOt, and
SiO2) were analyzed using a Thermo-ARL Advant’XP+
sequential X-ray fluorescence spectrometer (XRF). Sample preparation,
analytical conditions, and precisions/accuracy of the analyses follow
procedures detailed in Johnson et al. (1999). The whole-rock major
element concentrations of Isua ultramafic samples were determined at the
State Key Laboratory for Mineral Deposit Research in Nanjing University,
China. Small fresh rock pieces of Isua ultramafic samples were firstly
crushed into gravel-size chips. Clean chips were then powdered to 200
mesh for major element analysis. Measurements of whole-rock major
elements were performed by using a Thermo Scientific ARL 9900 XRF. The
measured values of diverse rock reference materials (BHVO-2 and BCR-2)
indicate that the uncertainties are less than ±3% for elements Si, Ti,
Al, Fe, Mn, Mg, Ca, K and P and less than ±6% for Na.
Trace element concentrations of Pilbara ultramafic samples were acquired
using an Agilent 7700 inductively coupled plasma mass spectrometer
(ICP-MS) in the Peter Hooper GeoAnalytical Laboratory at Washington
State University. Sample preparation, analytical conditions, and
precisions/accuracy of the analyses can be found in detail in Knaack et
al. (1994). Trace element contents of Isua ultramafic samples were
obtained at the University of Leeds, the U.K. These Isua ultramafic
samples were first crushed into powders with a ball mill. Details of
sample preparation, analytical procedures, and precisions are as
follows. First, whole-rock major elements were obtained using a Thermo
Scientific ARL 9900 X-ray fluorescence spectrometer. The measured values
of rock reference materials BHVO-2 and BCR-2 indicate that the
uncertainties on major element abundances are less than ±3 % for Si,
Ti, Al, Fe, Mn, Mg, Ca, K, and P and less than ±6 % for Na. Second, for
trace element analyses, about 100 µg sample powders and the reference
materials BHVO-1 and JP1 were digested and dissolved with
HNO3, HCl and/or HF and diluted with ultrapure water to
give a 1000-fold dilution in 3% HNO3. Samples were
analyzed for their trace element content (Sc, Ti, V, Cr, Mn, Co, Ni, Cu,
Sr, Ba, Th, U, Zr, Rb, and rare earth elements) using a Thermo
Scientific ICapQc ICP–MS at the University of Leeds. All concentrations
were corrected for uncertainties associated with weighing and diluting
the samples to produce a 1000-fold dilution. Reproducibility of the
BHVO-1 reference material during the analyses was ± 6 % for the rare
earth elements, and better than ± 15 % for all other elements, with the
exception of V and Th (± 16 and 18 %, respectively). Reproduction of
transition metal concentrations in JP1 was better than 10 relative %
for all elements (Sr, Ba, Th, U, and the rare earth elements were below
the detection limit of the instrument).
The major element contents of spinel crystals in Pilbara ultramafic
samples were obtained in situ from petrographic thin sections using a
JEOL JXA8230 Electron Probe Microanalyser (EMPA) at the University of
Leeds, U.K. Major element mineral (e.g., olivine, spinel, and
serpentine) compositions of the Isua ultramafic samples were analyzed in
situ on petrographic thin sections by a JEOL JX8100 Electron Probe
Microanalyser at the Guangzhou Institute of Geochemistry, Chinese
Academy of Sciences. At the same facility, a Carl Zeiss SUPRA55SAPPHIR
Field Emission Scanning Electron Microscope was used to collect images
of the Isua ultramafic samples.
The HSE concentrations and Re–Os isotopic data were obtained at the
Institute of Geology of the Czech Academy of Sciences, Czech Republic,
using the methods detailed in Topuz et al. (2018). In brief, the samples
were dissolved and equilibrated with mixed185Re-190Os and191Ir–99Ru–105Pd–194Pt
spikes using Carius Tubes (Shirey and Walker, 1995) and reverse aqua
regia (9 ml) for at least 72 hours. Decomposition was followed by Os
separation through solvent extraction by CHCl3 (Cohen
and Waters, 1996) and Os microdistillation (Birck et al., 1997).
Iridium, Ru, Pt, Pd, and Re were separated from the remaining solution
using anion exchange chromatography and then analyzed using a sector
field ICP-MS Element 2 (Thermo) coupled with Aridus
IITM (CETAC) desolvating nebulizer. The isotopic
fractionation was corrected using a linear law and standard Ir, Ru, Pd,
Pt (E-pond), and Re (NIST 3143) solutions that were run with samples.
In-run precision of measured isotopic ratios was always better than
±0.4% (2 σ). Os concentrations and isotopic ratios were obtained using
negative thermal ionization mass spectrometry (Creaser et al., 1991;
Völkening et al., 1991). Samples were loaded with concentrated HBr onto
Pt filaments with Ba(OH)2 activator and analyzed as
OsO3- using a Thermo Triton thermal
ionization spectrometer with Faraday cups in dynamic mode, or using a
secondary electron multiplier in a peak hopping mode for samples with
low Os concentrations. Internal precision for187Os/188Os determination was always
equal to or better than ±0.2% (2 σ). The measured Os isotopic ratios
were corrected offline for oxygen isobaric interferences, spike
contribution and instrumental mass fractionation using192Os/188Os = 3.08271 (Shirey and
Walker, 1998).
Literature data of Isua ultramafic rocks, crustal cumulates, and mantle
peridotites are compiled for comparison (see figure captions for data
sources). Fe contents of all complied data were recalculated to
represent FeOt using the procedure in Gale et al. (2013). Results were
plotted with GCDKit freeware developed by
Janoušek et al. (2006).
Results
Petrographic observations
We performed thin section petrographic analysis of both Isua and Pilbara
ultramafic samples to observe rock microtextures and mineral assemblages
that reflect igneous and metamorphic signatures, insofar as these
signatures are not obscured by alteration. Isua ultramafic samples show
varying degrees of alteration (Fig. 2, Fig. S1 ). Samples
AW17724-1, AW17724-4, and AW17725-2B are dominated by serpentine,
magnetite and carbonate minerals. Olivine, pyroxene, or protolith
textures can not be observed in these samples (Fig. 2b, 2d, Fig.
S1 ). Olivine grains are preserved in three samples (i.e., AW17724-2C
from lens B, AW17725-4 from lens B and AW17806-1; Fig. 2a, 2c ).
In sample AW17724-2C, Ti-humite phases (Ti-Clinohumite/Ti-Chondrodite)
are found as coexisting with olivine, serpentine and magnesite. In
addition, Ti-humite phases also occur as individual grains in the
matrix. Magnesite is sporadically found with Ti-humite phases and
olivine (Fig. 2a ). Although some Ti-humite, magnesite and
magnetite occur as apparent inclusions in olivine of AW17724-2C, they
are typically associated with cracks facilitating effective element
exchange between the central part of the olivine and the matrix
(Fig. 2a ). Notably, sample AW17724-2C also preserves a minor
volume of talc. Retrograde alteration in this sample is characterized by
younger lepidoblastic serpentine minerals cross-cutting or overgrowing
olivine, Ti-humite and magnesite (Fig. 2a ). In contrast,
olivine-bearing samples AW17725-4 and AW17806-1 do not exhibit any
Ti-humite phases. In addition to serpentinization, AW17725-4 shows
evidence of talc alteration, whereas sample AW17806-1 equally records
tremolite as an alteration product (Fig. 2c ). Relict olivine
grains preserved in sample AW17725-4 show polygonal textures, but the
protolith textures of AW17806-1 are altered beyond recognition
(Fig. 2b, 2d) .
In contrast to Isua samples, Pilbara samples have experienced complete
serpentinization and minor carbonation, such that no primary
ferromagnesian silicates can be identified (Fig. 3a–c, Fig.
S2 ). In all Pilbara samples, serpentine grains form clusters that show
similar extinction. Many such clusters have quasi-equant granular
outlines. We interpret these serpentine clusters to be pseudomorphs
after olivine. The interstitial space between the olivine-shaped
clusters is occupied by chlorites and/or Fe-Cr-Ti oxide minerals
(Fig. 3a–b ) or serpentine (Fig. 3a–c ). The
olivine-shaped serpentine clusters appear to form self-supporting
structures. Some interstitial serpentine clusters appear to preserve two
pairs of relict cleavages at ~90°, indicating a pyroxene
precursor (Fig. 3a ). Some interstitial serpentine clusters are
larger than the olivine-shaped serpentine clusters and enclose many of
the latter grains (illustrated in Fig. 3c : two sets of
serpentine clusters can be recognized via different brightness due to
extinction). Such patterns resemble poikilitic textures in which
early-formed chadacrysts are surrounded by younger, large oikocrysts
(Johannsen, 1931). In some locations, the olivine-shaped serpentine
clusters are compacted, forming polygonal textures (Fig. 3c ).
Late-stage alterations veins/cracks can be seen in samples AW52514-4A
and AL52614-4A (Fig. 3b ).