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. (1988a), and
the Geological Survey of Western
Australia 2013 database.
Geological background and proposed tectonic models
The Isua supracrustal belt
The ~35-km-long, ~1–4 km-wide Isua
supracrustal belt of southwestern Greenland is Earth’s largest
recognized Eoarchean terrane (Fig. 1a ). The protoliths of the
belt formed dominantly at ~3.8 Ga and
~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 and
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 (a high
Al2O3/TiO2 ”boninitic”
series and a low
Al2O3/TiO2“tholeiitic” series, Szilas et al. 2015), chert, banded iron
formation, and minor metamorphosed ultramafic igneous rocks, felsic
volcanic rocks, and detrital sedimentary rocks (e.g.,
Nutman et al., 2002;
Nutman and Friend, 2009).
Ultramafic rocks in the Isua area occur as ~1- to
~100-m-scale lenticular bodies associated with mafic
pillow lavas (e.g., Dymek et al.,
1988b; Szilas et al., 2015) and 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 (lens A in the south and lens B in the north)
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 (e.g., Friend et al. 2002;
Friend and Nutman, 2011; Nutman and Friend, 2009; Szilas et al., 2015).
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;
Guotana et al., 2022; Waterton et
al. 2022). These 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., 2022; 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 high primitive mantle-normalized Os, Ir
and Ru versus Pt and Pd preserved in ultramafic enclaves in the south
meta-tonalite (Van de Löcht et al., 2018), and similar patterns
preserved in the two meta-peridotite lenses of the Isua supracrustal
belt (Waterton et al. 2022).
The Isua supracrustal belt has been mostly interpreted to record
~3.8–3.6 Ga plate tectonic processes, including
subduction and subsequent extension (e.g.,
Arai et al., 2015; Nutman et al.,
2020; Nutman et al., 2013b; Nutman
and Friend, 2009). The presence of dunites in meta-peridotite lenses A
and B has been used to support such a plate tectonic origin (e.g.,
Friend and Nutman, 2011; Nutman et al., 2020; Van de Löcht et al. 2020)
as these dunites were interpreted as highly depleted mantle residues
tectonically thrust atop of supracrustal rocks in an Eoarchean
subduction setting (see Figure 8 of
Nutman et al., 2013a). In this
context, Isua dunites were interpreted as initially melt-depleted
olivine ± pyroxene ± spinel mantle residues. These residues experienced
fluid- and/or rock-dominated serpentinization, and UHP metamorphism, as
well as melt percolation in an Eoarchean mantle wedge, such that Isua
dunites preserve Ti-humite, variably fractionated HSE patterns, REE
patterns parallel to those of nearby basalts, and/or olivine with
clinopyroxene inclusions and mantle-like oxygen isotopes (e.g., Friend
and Nutman, 2011; Nutman et al.,
2020, 2021a). Olivine B-type fabrics were interpreted as recording
deformation in the mantle wedge (Kaczmerak et al., 2016). The deformed
and variably altered sub-arc mantle residues were then juxtaposed with
Isua supracrustal rocks via thrusting in an Eoarchean suprasubduction
zone (e.g., Nutman et al., 2020) and experienced additional modification
during and after Eoarchean (e.g., Nutman et al., 2021a; Guotana et al.
2022).
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 transport of materials, but the main driving force of
this transport is volcanic advection rather than gravitational
instability (Moore and Webb, 2013;
O’Reilly and Davies, 1981).
Voluminous mafic volcanism causes heat to be lost to the
atmosphere/space, and extensive volcanic depositional resurfacing as
well as burial and downwards advection of cold surface materials. The
volume loss from the ascent of hot magmatic materials is ultimately
balanced by the descent of the cold volcanic materials. At great depths,
portions of buried hydrated mafic crust are partially melted, forming
tonalitic melts. Other lower crustal rocks (along with varying fractions
of their fluid components) are recycled into the convecting mantle.
Therefore, in contrast to the idea that hot stagnant-lid regimes should
lack material exchange between surface and mantle (e.g.,
Nutman et al. 2021b), volcanic
advection in a heat-pipe setting is an efficient mechanism to generate
crust recycling and fluid-fluxing between crust and mantle (e.g., Moore
and Webb, 2013; O’Reilly and Davies, 1981). 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; Moore and Webb, 2013); or (2) contraction during a plate-breaking
and subduction event as or soon after the heat-pipe cooling ceases
(Beall et al., 2018; Moore and Webb, 2013; Tang et al. 2020).
Alternatively, deformation of a preserved fragment of heat-pipe
lithosphere may be possible at any subsequent time when involved in a
deformation zone of any tectonic setting. With respect to the formation
of ultramafic rocks, this model does not involve the thrusting of mantle
rocks atop crustal rocks, given that subduction and associated mantle
wedge settings do not occur during heat-pipe cooling. Therefore, the
heat-pipe model requires all Isua ultramafic rocks to represent high MgO
lavas (e.g., komatiites) or cumulates formed in magma chambers. In this
context, the geochemical signatures of Isua ultramafic rocks were
controlled by parental melt compositions, fractional crystallization,
melt-cumulate mixing and re-equilibration, and/or fluids/materials
released from crustal rocks. Their rock textures were produced by
crystallization of melts and/or subsequent deformation/mineral
re-equilibration under crustal conditions. Metamorphic assemblages
observed in Isua ultramafic rocks were formed under amphibolite facies
conditions, consistent with other parts of the belt (e.g.,
Ramírez-Salazar et al., 2021; Mueller et al., pre-print; cf. Friend and
Nutman, 2011; Nutman et al. 2020).
The East Pilbara Terrane
The ~40,000 km2 East Pilbara Terrane
of northwestern Australia is Earth’s largest and best-preserved
Paleoarchean terrane (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 layered ultramafic 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 have mostly occurred from
~3.5–3.2 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 (described
below).
Ultramafic rocks of the East Pilbara Terrane occur as layers or pods
interleaved with supracrustal rocks and 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 ~1–5 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. The East Pilbara Terrane exposes 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 ~3.53–3.43 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).
Most researchers interpret that 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) although a plate tectonic origin has also been proposed
(e.g., Kusky et al., 2021). One
subcategory of this tectonic regime is the partial convective overturn
cooling model (Collins et al., 1998), which 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. In
this model, (ultra)mafic magmatism associated with mantle plumes
produces 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 and folding of
supracrustal rocks deformed into synclines surround 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). Indeed, no lithology so far in the East Pilbara
Terrane has been interpreted as tectonically emplaced mantle rocks
(Hickman et al., 2021). Thus, Pilbara ultramafic rocks can be used as
non-plate tectonic crustal products to compare with Isua ultramafic
rocks.
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 (lens B
in the north) and AW17725-4 (lens A in the south) were collected from
the two meta-peridotite lenses which have been previously interpreted 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), Friend and Nutman (2011) and Waterton et al.
(2022) (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 all 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).
Modelled cumulates (Mallik et al.
2020) and variably depleted and refertilized mantle rocks (e.g.,
Chin et al. 2014, 2018) are also
compiled for comparison.
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
fragments 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 Nanjing
Hongchuang Exploration Technology Service Co., China. About 100 grams of
solid samples from each Isua ultramafic sample were first crushed into
smaller grains with a corundum jaw crusher. They were then crushed into
fine powder using an agate ball mill. Details of sample preparation,
analytical procedures, and precisions are as follows. The digestion
method of silicate rock samples is closed pressure acid dissolution
method. The specific steps are as follows: 50 mg of rock powder were
weighed directly into a steel-jacketed high-pressure polytetra
fluoroethylene bomb and then dissolved using an acid mixture of 1.5 mL
of 29 mol/L HF and 1 mL of 15 mol/L HNO3 at 190 °C for
72 hours. Then, the digested solution was evaporated to wet salt and
treated twice with 1 mL of concentrated HNO3 to avoid
the formation of fluorides. Finally, the evaporated residue was
dissolved with 1.5 mL HNO3 and 2 mL H2O
and the Teflon bomb was resealed and placed in the oven at 190 °C for 12
hours. The final solution was transferred to a polyethylene bottle and
diluted to 50 mL using H2O. Trace element analyses were
performed on an Agilent 7900 inductively coupled plasma mass
spectrometry (ICP-MS). The total quantitative analyses of trace elements
were achieved by external standard BCR-2 and BHVO-2 and internal
standard Rh dopped on line using an Agilent 7900 ICP-MS wet plasma. All
elements are repeatedly scanned for five times, which precision 1RSD are
better than 5 %. The margin of error of all trace element results for
rock powder reference materials was guaranteed to be plus or minus
within 10 %.
The major element compositions of spinel from the Pilbara ultramafic
samples were obtained using a JEOL JXA8230 Electron Probe Microanalyser
(EMPA) at the University of Leeds, United Kingdom. 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 Guangzhou 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 alteration signatures, as these are important
for the interpretation of geochemistry of altered samples. 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 carbonates with the absence of olivine,
pyroxene, or protolith textures (Fig. 2d, Fig. S1 ). On the
other hand, olivine grains are preserved in three samples (i.e.,
AW17724-2C from lens B, AW17725-4 from lens B and AW17806-1;Fig. 2a –c ), where they are cross-cut or overgrown by
retrograde serpentine minerals (Fig. 2a ). In addition to
serpentinization, meta-peridotite lens samples AW17724-4C and AW17725-4
show varying degrees of carbonitization (Fig. 2a–b ), whereas
sample AW17806-1 records tremolite as an alteration product
(Fig. 2c ). Small (submicron to ~20 μm)
serpentine, magnesite, and/or magnetite can be found within olivine
grains as inclusions or alteration products associated with cracks/veins
not visible on the thin section planes (Fig. 2a–b ). Relict
olivine grains preserved in sample AW17725-4 show polygonal textures,
but the protolith textures of AW17806-1 and AW17724-2C are altered
beyond recognition (Fig. 2b–c) . Ti-humite phases only occur in
AW17724-4 (Fig. 2a ; see Mueller et al. pre-print for detailed
petrological observations for this sample).
In contrast to Isua samples, Pilbara samples have experienced complete
serpentinization and minor carbonitization, 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 chlorite 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 ).