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. 2ac ), 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 ).