Igor S Puchtel
email: ipuchtel@umd.edu

Andrea Mundl-Petermeier

Mary Horan

Eero J Hanski

Janne Blichert-Toft

Richard J Walker

Ultra-depleted 2.05 Ga komatiites of Finnish Lapland: Products of grainy 1 late accretion or core-mantle interaction? 2 3 4 5

Keywords: Hf-W, Sm-Nd, Lu-Hf, Re-Os and Pt-Os isotope systems, highly siderophile 31 elements, Paleoproterozoic komatiites, grainy late accretion, core-mantle interaction, Finnish 32

Rhenium-Os, Pt-Os, Sm-Nd, Lu-Hf, and Hf-W isotope data, together with lithophile trace 36 element and highly siderophile element (HSE: Re, Os, Ir, Ru, Pt, and Pd) abundances, are 37 reported for 2.05 Ga Jeesiörova and Kevitsa komatiites from the Central Lapland Greenstone 38 Belt, Fennoscandia, Finland. Both komatiites are closely genetically related, with the Kevitsa 39 dikes having served as feeding magma conduits to the Jeesiörova pillowed and massive lavas. 40

The parental komatiite magma is estimated to have contained ~25 wt.% MgO and was, thus, 41 derived from a mantle source at least as hot as those of some of its late Archean counterparts.

42

A suite of Jeesiörova and Kevitsa whole-rock komatiite samples and olivine and chromite 43 separates define an internal Re-Os isochron with an age of 2049±13 Ma and an initial  187 Os 44 = -0.2±0.2 (2SE), indicating long-term chondritic Re/Os in the mantle source. By contrast, Pt-45 Os data for a set of Jeesiörova chromite separates define an average initial  186 Os = +29±2 46 (2SE), indicating a long-term history of suprachondritic Pt/Os in the mantle source. The 47 absolute HSE abundances in the mantle source of the Jeesiörova-Kevitsa komatiite system are 48 estimated to have been 120±5% of the present-day Bulk Silicate Earth (BSE). This is the first 49 komatiite system for which excess HSE in the mantle source, relative to modern BSE, has 50 been documented. 51 The 147 Sm-143 Nd and 176 Lu-176 Hf data yield isochron ages and initial ratios of, 52 respectively, 2046±22 Ma with  143 Nd = +3.7±0.3, and 2072±20 Ma with  176 Hf = +8.7±0.4 53 (2SE), indicating a long-term history of depletions of Nd relative to Sm, and Hf relative to Lu. 54 The measured  182 W = +1.5±3.3 is indistinguishable from the modern mantle value. Despite 55 being strongly depleted in highly incompatible lithophile trace elements, the Th-Nb-La 56 systematics of the komatiites indicate ~1% crustal contamination of the original komatiite 57 magma, assuming the contaminant was similar in composition to the calculated 58 Fennoscandian Tonalite Average (FTA). This level of contamination would have also 59 significantly modified the Nd, Hf, and W isotope compositions of the original komatiitic 60 magma, but not the Os isotope compositions or HSE abundances. The calculated original 61 komatiite magma, corrected for the effects of crustal contamination, would have had initial 62  143 Nd ~ +4.9,  176 Hf ~ +10.2, and  182 W ~ -10. 63 Our modeling indicates that the initial 186,187 Os/ 188 Os isotopic compositions and 64 suprachondritic HSE abundances, coupled with the projected negative  182 W, are best 65 explained by either (1) derivation from a mantle domain characterized by an excess of late 66 accreted, differentiated planetesimal core metal, i.e., "grainy" late accretion, or (2) addition of 67 chemically fractionated terrestrial core metal to the mantle source domain of the komatiites. 68

The presence of these characteristics in the Jeesiörova-Kevitsa komatiite mantle source 69 provides further evidence for the early creation and long-term survival of chemically diverse 70 domains within the mantle. 71 72 komatiitic melts are much more similar to the initial mantle source compositions than those in 107 basaltic melts. These high degrees of partial melting also led to comparatively high Os 108 abundances in komatiitic melts, which make their Os isotopic compositions insensitive to 109 crustal contamination. The high Os concentrations in komatiites, coupled with typically low 110 Re/Os ratios, mean that they can provide precise initial Os isotopic compositions of mantle 111 sources. 112 Compared to Archean komatiites, HSE and Os isotopes have been much less extensively 113 studied in their Proterozoic counterparts, primarily due to the paucity of Proterozoic 114 komatiites in the rock record. Here, we report combined 147 Sm-143 Nd, 176 Lu-176 Hf, 187 Re-115 187 Os, 190 Pt-186 Os, 182 Hf-182 W, and HSE and lithophile trace element abundance data for the 116 Proterozoic Kevitsa komatiitic dikes and associated Jeesiörova komatiites, all part of the 117

 120 Huhma et al., 2018). They hence provide ample opportunity for dating and obtaining 121 information on the composition of the original komatiitic magmas. We use these data to (i) 122 constrain the long-term evolution of the mantle domain from which the Fennoscandian Shield 123 was produced and which also gave rise to the komatiite parental magmas, and (ii) further 124 evaluate the extent of Archean-Paleoproterozoic mantle heterogeneity. 125

Geological background 126

. 130

. Stratigraphically, the 135 lower komatiites can be correlated with the komatiitic lavas of the Vetreny Belt in the SW 136 part of the Fennoscandian Shield, which have an internal Re-Os isochron age of 2407±6 Ma 137 (Puchtel et al., 2016b). 138 Mineral compositions were determined using a JEOL JXA-8530F Plus Hyper Probe instrument at the 204 Center of Material Analysis, University of Oulu. The analytical conditions were an accelerating voltage of 15 kV 205 and a beam current of 15 nA. Peak and background counting times were 10 and 5 s, respectively. The standard 206 built-in ZAF correction routine was used.

).

218

The average total analytical blank (TAB) measured during the present analytical campaign was (in pg): Ru 219 6.0, Pd 17, Re 0.53, Os 0.42, Ir 1.2, and Pt 174 (N = 9). For the whole-rock komatiite samples, the average TAB 220 constituted less than 0.1% for Os, Re, Ir, Ru, and Pd, and less than 1% for Pt of the total amount of element 221 analyzed. For the olivine and chromite separates, the TAB for Os and Ir constituted less than 0.1%, for Ru less 222 than 0.4%, for Re between 0.1 and 3%, for Pt between 14 and 19%, and for Pd less than 1% of the total amount 223 of element analyzed. We, therefore, cite ±0.1% as the uncertainty on the concentrations of Os, ±0.2% for Ir, 224 between 0.3 and 0.4% for Ru, between 0.1 and 1% for Pd, between 0.1 and 3% for Re, and between 1.0 and 19% 225 for Pt, of the total amount of element analyzed. The uncertainty on the Re/Os ratio for each sample was 226 calculated by multiplying the estimated uncertainties on the Re and Os abundances. These uncertainties vary 227 between 0.30 and 3.0% relative.

228 4.3. Pt-Os isotopic data.

229

In order to obtain the amount of Os required for high-precision measurements of 186 Os/ 188 Os and 187 Os/ 188 Os 230 (~100 ng), ~1.0 g of pure chromite separates from Jeesiörova komatiite samples 12D-PPR and 13-EJH, 231 representing different fractions in terms of magnetic susceptibility and density, were digested in Carius tubes.

232

For the initial unspiked digestions, six ~1 g fractions of chromite separate 12D-PPR and two fractions of 13-EJH 233 were digested. After the digestion was complete, the tubes were chilled, opened, and ~2% of the acid sample 234 solution from each CT were transferred into a 25 mL Pyrex™ CT for precise determination of Pt/Os, Re/Os, and 235 Ir/Os. The spiked aliquots were processed using the same procedure utilized for the Re-Os and HSE analyses.

236

From the remaining part of the unspiked acid sample solutions, Os was extracted and purified using the same 237 protocol utilized for the Re-Os work. The Os cuts from the batch of two CT containing sample 13-EJH were 238 combined into one cut and used for the precise measurements of 186 Os/ 188 Os and 187 Os/ 188 Os in this sample. For 239 sample 12D-PPR, every Os cut from six individual CT was used for the precise measurements of Os isotopic 240 compositions. For TAB-corrections of the data from the spiked aliquots, since the ID digestion introduced ~98% 241 of the total blank, the TAB of the initial (IC) digestion was considered negligible. 242 8 The high-precision measurements of the 186 Os/ 188 Os and 187 Os/ 188 Os ratios were performed by N-TIMS in 243 static mode on a ThermoFisher Triton ® mass spectrometer at the IGL. During each run, between 1200 and 2400 244 ratios were collected for each sample load; the in-run uncertainties on the measured 186 Os/ 188 Os and 187 Os/ 188 Os 245 ratios are quoted as 2SE. The possible isobaric interference of 186 W 16 O 3 -on 186 Os 16 O 3 -was monitored and 246 corrected for by measuring masses 184 OsO 3 -and 183 W 16 O 3 -using the electron multiplier. 247 The mean of the Johnson-Matthey Os standard runs during the period of data collection was 0.001302±2 for 248 184 Os/ 188 Os, 0.1198432±18 (±15 ppm) for 186 Os/ 188 Os, and 0.1137950±18 (16 ppm) for 187 Os/ 188 Os (2SD, N = 249 14); these long-term reproducibilities were used to assess the true uncertainty on the measured 186 Os/ 188 Os and 250 187 Os/ 188 Os ratios for individual samples.

251 4.4. Tungsten isotopic compositions and abundances 252 The W isotope analyses were carried out at the IGL. For each sample, approximately 100 grams of powder 253 were processed to obtain the ~1 g of W necessary for high-precision W isotope measurements. The sample 254 powders for each sample were digested in four 300 mL Savillex Teflon screw-cap vials using a 5:1 mixture of 255 double-distilled concentrated HF and HNO 3 on a hot plate at 150°C for one week. Tungsten was separated and 256 purified using the four-stage ion-exchange chromatography protocol described in Peters et al. (2019), with minor 257 modifications. The third stage involving a 1.5 mL anion-exchange column was repeated to improve the 258 separation of Ti from W, and this step significantly increased W ionization efficiency. Tungsten recovery using 259 this procedure was better than 90%. 260 Tungsten isotopic compositions were measured by N-TIMS on a ThermoFisher Triton mass spectrometer at 261 the IGL using a 2-line multi-static acquisition protocol and following a slightly modified technique described by 262 Archer et al. (2017). All data are reported as µ 182 W and µ 183 W, which are the part per million (ppm) deviations 263 of 182 W/ 184 W and 183 W/ 184 W, respectively, in a given sample from those of the in-house Alfa Aesar laboratory W 264 standard. Uncertainties on  182 W, based on the long-term 2SD of our Alfa Aesar laboratory standard, were ±4.5 265 ppm. The measured 183 W/ 184 W ratios were identical within uncertainties (±6  183 W units) to those in the average 266 Alfa Aesar standard data. 267 Tungsten abundances were determined by isotope dilution ICP-MS. Tungsten was purified using a 268 previously established anion-exchange chromatography technique (e.g., Kleine et al., 2004). Concentrations 269 were measured using the Element 2 single-collector ICP-MS at the PL. Typical uncertainties were ~5% relative.

270 4.5. Sm

Introduction 73

The chemically and isotopically heterogeneous nature of the mantle has long been 74

established, largely through the study of modern rocks (e.g., [START_REF] Gast | Isotopic composition of lead and strontium from 1144 Ascension and Gough islands[END_REF]Hart and 75 Brooks, 1977;[START_REF] Zindler | Nd, Sr and Pb isotopic systematics in a three-1480 component mantle: a new perspective[END_REF][START_REF] White | Sr and Nd isotope geochemistry of oceanic basalts 1460 and mantle evolution[END_REF][START_REF] White | Hf-Nd-Sr isotopes and incompatible element abundances 1462 in island arcs: implications for magma origins and crust-mantle evolution[END_REF]76 Hofmann, 198476 Hofmann, , 1997;;[START_REF] Jacobsen | Isotopic and chemical constraints on mantle-crust evolution[END_REF][START_REF] Galer | Early mantle differentiation and its thermal 1133 consequences[END_REF]. However, the origin of 77 the heterogeneities for different elements and isotopic systems (e.g., 146,147 Sm-142,143 Nd, 176 Lu-78 176 Hf, 182 Hf-182 W), as well as their length scales and residence times, remain topics of debate. 79 Some of the heterogeneities have been argued to be primordial, reflecting initial planetary 80 accretion/differentiation, including magma ocean crystallization processes (e.g., Goldstein 81 and Galer, 1992;[START_REF] Albarède | Hf-Nd 1031 isotope evidence for a transient dynamic regime in the early terrestrial mantle[END_REF][START_REF] Boyet | 142 Nd evidence for early (> 4.53 Ga) global differentiation 1080 of the silicate Earth[END_REF][START_REF] Touboul | 182 W Evidence for Long-Term Preservation 1427 of Early Mantle Differentiation Products[END_REF]82 Jacobsen and Yu, 2015;[START_REF] Puchtel | Insights into early Earth from Barberton komatiites: Evidence 1355 from lithophile isotope and trace element systematics[END_REF]2016a;Rizo et al., 2016b;[START_REF] Mundl | Tungsten-182 heterogeneity in modern ocean island basalts[END_REF]83 Mundl-Petermeier et al., 2019;[START_REF] Tusch | Uniform 182 W isotope compositions in Eoarchean rocks from the Isua 1436 region, SW Greenland: The role of early silicate differentiation and missing late veneer[END_REF], while others may have originated as a 84 result of a protracted accretion history [START_REF] Willbold | The tungsten isotopic composition of the Earth's 1465 mantle before the terminal bombardment[END_REF][START_REF] Kruijer | Lunar tungsten isotopic 1233 evidence for the late veneer[END_REF]Touboul et 85 al., 2015;[START_REF] Puchtel | 182 W and HSE 1368 constraints from 2.7 Ga komatiites on the heterogeneous nature of the Archean mantle[END_REF][START_REF] Archer | 1045 Lack of late-accreted material as the origin of 182 W excesses in the Archean mantle: 1046 Evidence from the Pilbara Craton, Western Australia[END_REF]. 86 Osmium isotope and highly siderophile element (HSE, including Re, Os, Ir,Ru,Pt,and 87 Pd) abundance systematics of the mantle through time provide information about certain 88 aspects of early Earth processes. For example, the observation that the HSE occur in 89 approximately chondritic relative proportions in the Bulk Silicate Earth (BSE), and that 90 absolute abundances of at least some of the HSE are higher than would be expected from 91 metal-silicate equilibration during core formation, have led to the concept of late accretion, 92 which purports that 0.5-1% of Earth's mass was accreted to the mantle subsequent to 93 cessation of core formation [START_REF] Kimura | Distribution of gold and rhenium between nickel-1208 iron and silicate melts; implications for abundance of siderophile elements on the Earth 1209 and Moon[END_REF][START_REF] Chou | Siderophile trace elements in the Earth's 1098 oceanic crust and upper mantle[END_REF][START_REF] Morgan | Osmium isotope constraints on Earth's late accretionary history[END_REF][START_REF] Morgan | Ultramafic xenoliths: Clues to Earth's late accretionary history[END_REF]. 94

Issues related to the nature of late accretion include the composition of the late accreted 95 materials and the time frame within which they were delivered to Earth and homogenized 96 within the mantle (e.g., Bennett et al., 2002;[START_REF] Maier | Progressive mixing of meteoritic veneer into the early Earth's deep 1252 mantle[END_REF]Coggon et al., 2013;Walker et 97 al., 2015;van de Löcht et al., 2018). 98

The Os isotopic composition and absolute HSE abundances of portions of the Archean 99 mantle have been the subject of considerable prior study (e.g., [START_REF] Brügmann | Noble metal 1095 abundances in komatiite suites from Alexo, Ontario, and Gorgona Island, Colombia[END_REF]Walker 100 et al., 1988;[START_REF] Foster | Re-Os isotopic evidence for 1130 genesis of Archaean nickel ores from uncontaminated komatiites[END_REF][START_REF] Gangopadhyay | Re-Os systematics of the ca. 2.7 Ga komatiites 1135 from Alexo, Ontario, Canada[END_REF][START_REF] Puchtel | Insights into 1358 Early Earth from the Pt-Re-Os isotope and Highly Siderophile Element abundance 1359 systematics of Barberton komatiites[END_REF]Puchtel et al., , 101 2016a,b;,b;2018;[START_REF] Maier | Progressive mixing of meteoritic veneer into the early Earth's deep 1252 mantle[END_REF]. These studies have mostly focused on komatiites because 102 komatiites are valuable probes of absolute HSE abundances and Os isotopic compositions in 103 some portions of the mantle. This stems from the fact that most komatiites were formed by 104 sufficiently high degrees of partial melting that the dominant sulfide hosts of HSE in the 105 mantle were dissolved into the melts. Under such conditions, the HSE abundances in 106

The upper komatiites are associated with phyllites and black schists of the Savukoski 139 Group whose sedimentary precursors were deposited on cratonic siliciclastic sediments of the 140 Sodankylä Group [START_REF] Hanski | Central Lapland Greenstone Belt[END_REF]. The upper komatiitic volcanic rocks extend 141 from Finnish Lapland to northern Norway over a distance of ~400 km and are associated with 142 Ti-rich picrites and basalts [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF]. Unlike typical Archean komatiites, the upper 143 komatiites are characterized by the abundance of volcaniclastic eruptions [START_REF] Saverikko | The pyroclastic komatiite complex at Sattasvaara in northern Finland[END_REF]; 144 [START_REF] Barnes | Ti-rich komatiites from Northern Norway[END_REF]. They also occur as massive lava flows, pillow lavas, tuffs, and rare 145 layered flows. The volcanic structures and types of the associated sedimentary rocks attest to 146 submarine conditions of eruption in a deepened sedimentary basin. Although belonging to the 147 Al-undepleted, or Munro-type of lavas of [START_REF] Nesbitt | Komatiites: geochemistry and genesis[END_REF] typical of late Archean 148 komatiite sequences, the upper komatiites are further distinct from Archean komatiites in 149 having elevated TiO 2 contents and, therefore, they were classified as a Ti-enriched komatiite 150 type [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF]. 151

The upper komatiites are best preserved in the Jeesiörova area studied by Hanski et al. 152 (2001). Magmatic olivine has not been reported to have been preserved in komatiitic lavas 153 anywhere in Finnish Lapland, but magmatic clinopyroxene is well preserved at Jeesiörova, 154 occurring as poikilitic intergrowths in olivine cumulates, as prismatic grains in gabbroic 155 rocks, and as needle-like crystals in pillowed and massive lavas. The presence of magmatic 156 clinopyroxene in the lavas allowed [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF] to obtain a weighted average Sm-Nd 157 isochron age of 2056±25 Ma for clinopyroxene-whole-rock pairs for the Jeesiörova 158 komatiites. The initial  143 Nd values range from +2 to +4 for individual samples, with the 159 most LREE-depleted samples having the highest initial  143 Nd values. This inverse correlation 160 between the La/Sm and  143 Nd values was interpreted to be the result of contamination by 161 material of the upper crustal rocks through which the Jeesiörova komatiite magma ascended 162 [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF]. 163 [START_REF] Gangopadhyay | Origin of Paleoproterozoic 1137 Ti-enriched komatiitic rocks from Jeesiörova, Kittilä Greenstone Complex, Finnish 1138 Lapland[END_REF] analyzed a set of whole-rock Jeesiörova komatiite samples 164 and chromite separates for Re-Os isotope systematics. Although the Re-Os system in the 165 whole-rock samples was found to be disturbed by post-magmatic processes, by analyzing 166 chromite separates, the authors constrained the initial 187 Os/ 188 Os ratio for the komatiite lavas 167 to be chondritic ( 187 Os(T) = +0.1±0.6). 168

The Kevitsa komatiitic dikes were discovered by [START_REF] Mutanen | Geology and ore petrology of the Akanvaara and Koitelainen mafic 1290 layered intrusions and the Keivitsa-Satovaara layered complex, northern Finland[END_REF] during an exploration 169 drilling campaign of the Kevitsa Ni-Cu sulfide deposit. The dikes were penetrated by at least 170 a dozen drill holes in an area of ~1 km 2 , covering the ore-bearing part of the Kevitsa mafic-171 ultramafic intrusion (Fig. 1c). This intrusion is located approximately 50 km east of the 172 Jeesiörova area within pelitic metasediments of the Savukoski Group (Fig. 1a) and has a U-173 Pb zircon age of 2058±4 Ma [START_REF] Mutanen | U-Pb geochronology of the Koitelainen, Akanvaara and 1296 Keivitsa layered intrusions and related rocks[END_REF], which is identical, within 174 uncertainty, to the Sm-Nd age of the Jeesiörova komatiites. Each of the drill cores contains 175 from one to four dike intersections at depths varying from 12 to 870 m beneath the surface. 176

Although the true thickness of the dikes is less than 7 m, the lengths of the intersections reach 177 up to 16 m because the drill holes cut through the dikes at sharp angles (≤30°). 178

The dikes are generally subvertical, with the dip varying between 65° and 90°, and strike 179 approximately ENE [START_REF] Mutanen | The Akanvaara intrusion and the Keivitsa-Satovaara Complex, with stops 1293 at Kaikkivaltiaanlehto and Särkivaara intrusions[END_REF]. They are characterized by strong internal 180 differentiation related to olivine fractionation, having an olivine-enriched central part and 181 finer-grained chilled zones near contacts with the wall rocks, which are composed of olivine-182 pyroxene cumulates of the Kevitsa intrusion. The differentiation is rather symmetrical across 183 the dikes, suggesting that originally they were also steeply dipping. The dikes are chemically 184 similar to the Jeesiörova komatiites and have been interpreted to represent their feeder magma 185 conduits [START_REF] Huhma | Sm-Nd and 1185 U-Pb isotope geochemistry of the Palaeoproterozoic mafic magmatism in eastern and 1186 northern Finland[END_REF]; the absolute age of the dikes, however, has so far remained 186 unknown. 187

Samples 188

Sampling of the Kevitsa komatiite dikes was done at the National Drill Core Depot of the 189 Geological Survey of Finland, Loppi, Finland. Samples for chemical and isotopic analyses 190 were selected from drill core DDH 814, which contains two dike intersections at depths of 191 152.10-159.55 and 175.00-188.50 m, with the corresponding estimated thicknesses of the 192 dikes being ~1.3 and ~6.8 m. 193 For this study, we also prepared powders from hand specimens of the best preserved and 194 most primitive komatiite lava samples 12D-PPR, 13-EJH, and 17.1-PPR from the Jeesiörova 195 area [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF]. In addition, we powdered sample LP-10 that was recently collected 196 from an olivine-phyric massive komatiite lava flow from the same area. 197 198

Analytical techniques 199

Details of the majority of the analytical techniques used in this study have been reported in a number of 200 previous publications (e.g., Puchtel et al., 2016a,b;2018); these, therefore, are only briefly summarized here. In 201 contrast, full details are provided for those techniques implemented here for the first time. precisely for the AMES Nd standard on the IGL ThermoFisher Triton (0.512152±2; [START_REF] Puchtel | 182 W and HSE 1368 constraints from 2.7 Ga komatiites on the heterogeneous nature of the Archean mantle[END_REF].

286

In order to assess the accuracy of the analyses, several separate powder aliquots of USGS GRM BIR-1 and 287 BCR-1 were processed and analyzed using the same analytical protocol. The measured 143 Nd/ 144 Nd ratios for 288 both GRM are identical, within their respective uncertainties, to the GeoRem preferred values (Supplementary 289 Table A2), as well as to the average high-precision value for BCR-1 obtained on the IGL ThermoFisher Triton 290 (0.512645±1; [START_REF] Puchtel | 182 W and HSE 1368 constraints from 2.7 Ga komatiites on the heterogeneous nature of the Archean mantle[END_REF].

291

Measurements of the Sm isotopic compositions were also performed on Faraday cups of the ThermoFisher 292 Neptune Plus ICP-MS at the PL in static mode. For each sample, 20 ratios were collected with 8 s integration 293 times in blocks of 10 ratios each. The effects of instrumental mass fractionation were corrected for relative to 294 147 Sm/ 152 Sm = 0.56081 using an exponential law. The uncertainty on the Sm/Nd ratio in the samples analyzed 295 was estimated on the basis of replicate analyses of the USGS GRM BCR-1 ( 147 Sm/ 144 Nd = 0.13941±0.00014) to 296 be 0.10% relative (2SD). 303 for one week. The solutions were dried down, 0.7 mL of concentrated SeaStar HClO 4 added to the residue, the 304 inserts re-sealed and kept on a hotplate at 180°C for 24 hours. The solutions were then dried down at ~230°C, 305 and the residues converted into the chloride form using 6M HCl. This step was repeated several times until clear 306 solutions were obtained, and then dried down.

307

At DTM, the residue was re-dissolved in 5 ml of a 1M HCl -0.1M HF mixture, and loaded onto a 0.6×20 308 cm column filled with AG50W-X8 200-400 mesh cation-exchange resin. High field strength elements, including 309 Hf, were eluted in the loading solution and an additional 5 mL of 1M HCl -0.1M HF. After eluting 44 ml of 310 2.5M HCl, heavy REE, including Lu, were eluted in 12 mL of 4M HCl [START_REF] Blichert-Toft | On the Lu-Hf isotope geochemistry of silicate rocks[END_REF]. Hafnium was 311 purified on a 0.6×10 cm column of Eichrom LN resin (100-150 m bead size) by loading the sample in 5 mL of 312 2.5N HCl, followed by washes of 10 mL 2.5M HCl, 10 mL 6M HCl, 4 mL H 2 O, 40 mL of a mixture consisting 313 of 0.09M Citric acid -0.45M HNO 3 -1% H 2 O 2 , 5 mL of the same solution without H 2 O 2 , and, finally, 20 mL of 314 6M HCl -0.06N HF. Hafnium was then eluted in 8 mL of 6M HCl -0.4M HF. Lutetium was purified using 315 Eichrom LN resin (50-100 m bead size) on a 0.4×7 cm column. The HREE fraction from the first column was 316 loaded in 2.5M HCl, followed by a wash of 30 mL 2.5M HCl to remove much of the other HREE, then Lu was 317 collected in 10 mL 6M HCl. The resultant Hf and Lu cuts were used for the measurements of the Hf and Lu 318 isotopic compositions.

319

Measurements of the Hf isotopic compositions were performed on the Faraday cups of the ThermoFisher 320 Neptune Plus ICP-MS at the PL in static mode followed by off-line correction for spike contributions. For each 321 sample, 200 to 300 ratios were collected with 8 s integration time in blocks of 100 ratios each. Before each 322 sample measurement, a 60 s baseline measurement was performed for each Faraday cup/amplifier pair without 323 beam deflection, and this baseline was then automatically subtracted from the sample beam. Instrumental mass 324 fractionation was corrected for relative to 179 Hf/ 177 Hf = 0.7325 using an exponential law. The Yb, Lu, Ta, and W 325 isobaric interferences were corrected for using 173 Yb/ 176 Yb = 1.256, 175 Lu/ 176 Lu = 37.70, 181 Ta/ 180 Ta = 8129, and 326 183 W/ 180 W = 109.0. A total of five 30 ppb JMC-475 Hf standard solutions were run at the beginning and end of 327 the analytical session, with 200 ratios collected during each measurement. During the measurements, the signal 328 intensities for both the standard and the samples were kept at constant levels, between 1.8 and 2.6 V on the 178 Hf 329 mass; the in-run precision of the measured 176 Hf/ 177 Hf ratio for both samples and standards was between 11 and 330 13 ppm (2SE). During the course of the present analytical campaign, the external reproducibility of the 30 ppb 331 JMC-475 Hf standard solution measurements for 176 Hf/ 177 Hf was ±14 ppm (2SD); this value was used to 332 estimate the true uncertainty on the Hf isotopic analyses. The average 176 Hf/ 177 Hf ratio measured for the JMC-333 475 Hf standard solution during the analytical session was used to calculate the instrumental mass bias 334 coefficient for correction of the measured 176 Hf/ 177 Hf ratio in the samples using the true value of 176 Hf/ 177 Hf = 335 0.282163±0.000009 as determined by a multiple dynamic analysis protocol of [START_REF] Blichert-Toft | Separation of Hf and Lu for high-1070 precision isotope analysis of rock samples by magnetic sector multiple collector ICP-MS[END_REF].

336

In order to assess the accuracy of the analyses, an aliquot of USGS GRM BCR-1 was processed and 337 analyzed using the same protocol as that used for the samples. The measured 176 Hf/ 177 Hf ratio of BCR-1 is 338 identical, within the uncertainty, to the GeoRem preferred value.

339

Measurements of the Lu isotopic compositions were also performed on the Faraday cups of the 340 ThermoFisher Neptune Plus ICP-MS at the PL in static mode. For each sample, 40 ratios were collected with 4 s 341 integration time in blocks of 20 ratios each. Before each sample measurement, a 60 s baseline measurement was 342 performed for each Faraday cup/amplifier pair without beam deflection, and this baseline was then automatically 343 subtracted from the sample beam. Instrumental mass fractionation was corrected for by analysis of a 10 ppb Lu 344 standard solution before and after every 4 samples. The Yb and Hf isobaric interferences were corrected for on-345 line using 176 Yb/ 173 Yb = 0.7962 and 176 Hf/ 177 Hf = 0.2822. The Lu concentration obtained for the SRM BCR-1 in 346 this study was identical, within the uncertainty, to the GeoRem recommended value. The propagated uncertainty 347 on the Lu/Hf ratio in the samples analyzed is estimated to be 0.5%. The total analytical blanks were 70 pg for Hf 348 and 2 pg for Lu. 349 350

Results 351

Petrography of the komatiites 352

The studied samples of the Jeesiörova komatiites are massive rocks with up to 1 mm-sized 353 euhedral, completely serpentinized olivine phenocrysts in a groundmass of prismatic or 354 needle-like clinopyroxene crystals and devitrified glass (Fig. 2a in Hanski and Kamenetsky, 355 2003). The rocks also contain euhedral chromite grains up to 0.3 mm in size, often containing 356 melt inclusions (Figs. 2b and 2d in Hanski and Kamenetsky, 2003). 357

Compared to the Jeesiörova rocks, the Kevitsa komatiites analyzed in this study are 358 characterized by a much better preserved magmatic mineralogy. The rocks are olivine-359 porphyritic, with the most coarse-grained varieties found in dike centers, due to the presence 360 of olivine phenocrysts (Fig. 2A,B,E). The size of the phenocrysts is commonly less than 6 361 mm in the center of the dikes, decreasing gradually to 0.1-0.2 mm near the dike margins. 362

Occasionally, large amoeboid and embayed megacrysts are observed, with a maximum grain 363 size of ~2 cm (Fig. 3). Olivine phenocrysts show various morphologies; they are usually 364 euhedral to subhedral, but sometimes have irregular shapes with ragged outlines (Fig. 2D). 365

The latter are interpreted to be broken pieces of larger phenocrysts. Also observed are 366 elongated olivine bars up to 1 cm in length, with an aspect ratio of up to 20 (Fig. 2B). 367

Although olivine phenocrysts contain ubiquitous chrome spinel inclusions 0.1-0.3 mm in size, 368 some of the largest phenocrysts are nearly devoid of them; in these cases, chrome spinel 369 grains are concentrated along the olivine grain boundaries. 370

The holocrystalline groundmass is composed of clinopyroxene, orthopyroxene, olivine, 371 plagioclase, brown amphibole, and opaque minerals, including two kinds of spinels, chrome 372 spinel, and lamellar ilmenomagnetite (Fig. 2C). These oxides occur as individual or 373 composite grains, in which chrome spinel is partly surrounded by ilmenomagnetite. 374

Plagioclase forms lath-like grains up to 1 cm in length, but other silicates are anhedral or 375 equant, with their grain size ranging commonly between 0.1 and 0.5 mm. 376

The chilled margins of the dikes contain olivine microphenocrysts (<0.5 mm in size) and 377 some chrome spinel grains in a fine-grained to cryptocrystalline groundmass composed 378 mostly of clinopyroxene and plagioclase. It is noteworthy that no vesicles are found in the 379 contact zones, indicating that either the pressure of emplacement was too high for volatiles to 380 form a gas phase, or the volatile content was very low. 381 382

Mineral chemistry 383

Representative electron probe analyses of igneous minerals from the Kevitsa komatiites 384 are listed in the Supplementary Table A3 Lapland, but distinct from the more Ti-rich spinels that occur in the associated picrites 403 [START_REF] Hanski | Chrome spinel-hosted melt inclusions in primitive 1159 Paleoproterozoic volcanic rocks, northern Finland: evidence for coexistence and mixing 1160 of komatiitic and picritic magmas[END_REF]. 404 5.2.4. Amphibole. The compositions of the brown groundmass amphibole (Fig. 2B,C) 405 straddle the boundary between Ti-rich pargasite and Ti-rich magnesio-hastingsite. The TiO 2 406 contents fall between 2.9 and 4.0 wt.%, consistent with the magmatic nature of the mineral. 407 Amphibole shows moderate Na 2 O contents of 2.3-3.0 wt.%, whereas K 2 O is low, 0.10-0.26 408 wt.%, consistent with the low level of K 2 O in the parental komatiite magma. 409

Major and lithophile trace element abundances 410

The major and trace element concentration data for the Jeesiörova and Kevitsa komatiites 411 are listed in Tables 1 and2, and selected elements are plotted against MgO contents on 412 variation diagrams in Fig. 4 and as BSE-normalized values in Fig. 5. 413

The MgO abundances range between 25.3 and 28.6 wt.% in the Jeesiörova komatiites and 414 between 16.2 and 28.3 wt.% in the Kevitsa komatiites. In the Kevitsa komatiite dikes, the 415 lowest abundances are observed in the chilled margin samples and the highest abundances in 416 the samples from the central parts. The abundances of TiO 2 , CaO, Al 2 O 3 and incompatible 417 lithophile trace elements in the dikes plot on well-constrained trends with negative slopes that 418 intersect the MgO axes at ~51 wt.%; these intersections correspond to the composition of 419 liquidus olivine obtained in this study and also reported for the dikes by [START_REF] Nicklas | Secular Mantle Oxidation across the Archean-Proterozoic 1310 Boundary: Evidence from V partitioning in Komatiites and Picrites[END_REF]. 420

The abundances of the compatible elements Ni and Cr show positive correlations with MgO 421 contents (Fig. 4). These strong correlations reflect olivine and chromite control during 422 magmatic differentiation of the dikes, and also indicate immobile behavior of the elements 423 analyzed during secondary alteration processes. 424

The four samples of the Jeesiörova komatiites analyzed in this study have slightly higher 425 Ti, Cr, La, Ce, Th, and U concentrations compared to the Kevitsa komatiites (Fig. 4), but for 426 the rest of major and trace elements, plot on the same trends. The Jeesiörova and Kevitsa 427 komatiites have Al 2 O 3 /TiO 2 (11.7±0.3 and 13.9±0.1, respectively, 2SE) that are lower, and 428 (Gd/Yb) N (1.53±0.08 and 1.49±0.01, respectively, 2SE) that are higher, than the BSE values 429 of 22.4 and 1.00, respectively. This is due to enrichment of the parental komatiite magma in 430 Ti and the middle REE, rather than to depletion in Al and the heavy REE, as is evident from 431 comparison with the typical Al-undepleted komatiites from Pyke Hill, Ontario, Canada (Fig. 432 4A). As such, they are classified as the Al-undepleted komatiite type of [START_REF] Nesbitt | Komatiites: geochemistry and genesis[END_REF]. 433

The BSE-normalized lithophile trace element abundances (Fig. 5) are characterized by 434 strong depletions in light REE ((La/Sm) N = 0.347±0.022 and 0.264±0.002 in the Jeesiörova 435 and Kevitsa komatiites, respectively; 2SE), sloping HREE ((Gd/Yb) N = 1.53±0.08 and 436 1.49±0.01, respectively; 2SE), and enrichments in Th and U relative to the neighboring 437 element with similar incompatibility, i.e., Nb ((Nb/Th) N = 0.595±0.019 and 0.583±0.011, 438 respectively). 439 440

Re-Os isotopic compositions and HSE abundances 441

The Re-Os isotopic and HSE abundance data for the Jeesiörova and Kevitsa komatiites 442 and olivine and chromite separates are listed in Table 3 andplotted Thirty-three whole-rock samples and olivine and chromite separates define a regression 446 line with a slope corresponding to an ISOPLOT Model 3 age of 2049±13 Ma and a chondritic 447 initial 187 Os/ 188 Os = 0.11284±28 ( 187 Os = -0.2±0.2). This age is in good agreement with the 448 internal Sm-Nd isochron age of 2056±25 Ma obtained by [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF] for the 449 Jeesiörova komatiites. The initial  187 Os = -0.2±0.2 is identical, within uncertainty, to the less 450 precise average initial  187 Os = +0.1±0.6 obtained for chromite separates from the Jeesiörova 451 komatiites by [START_REF] Gangopadhyay | Origin of Paleoproterozoic 1137 Ti-enriched komatiitic rocks from Jeesiörova, Kittilä Greenstone Complex, Finnish 1138 Lapland[END_REF]. 452

In CI chondrite-normalized HSE diagrams (Fig. 7), the Jeesiörova and Kevitsa komatiite 453 samples display typical komatiitic patterns with moderate enrichments in incompatible Pt, Pd,454 and Re relative to compatible Os, Ir, and Ru ((Pd/Ir) N = 2.84-10.1). By contrast, the two 455 olivine separates analyzed show an inverse relationship, being enriched in compatible relative 456 to incompatible HSE, with the exception of the elevated Re abundance in one of the separates. 457

This elevated Re concentration is, however, supported by the correspondingly more 458 radiogenic Os isotopic composition of this separate, indicating that the olivine likely 459 contained inclusions of interstitial melt enriched in incompatible elements. 460

In the MgO versus HSE diagrams (Fig. 8), variations in the Os and Ir contents show broad 461 positive correlations with the MgO abundances, with the large degree of scatter most likely 462 due to the presence of Os-Ir alloy on the liquidus, along with olivine. The Ru abundances 463 display a more regular behavior and were likely mostly controlled by chromite fractionation 464 during differentiation of the komatiites. 465

The Pt, Pd, and Re abundances show strong negative correlations with MgO contents and, 466 with the exception of some scatter in the Re content, especially among the Jeesiörova 467 komatiites, plot on the trends that pass through the olivine compositions; these variations are 468 consistent with olivine control, indicating both incompatible behavior of these HSE during 469 komatiite differentiation and their immobile behavior during post-magmatic processes. Of 470 note, the Jeesiörova and Kevitsa rocks show a subchondritic (Pt/Pd) N = 0.78±0.01, which is 471 typical of komatiites (e.g., Puchtel et al., 2009b), but is ~20% lower than the near-chondritic 472 ratios in the Ni-Cu-PGE sulfide deposits of Finnish Lapland [START_REF] Mutanen | Geology and ore petrology of the Akanvaara and Koitelainen mafic 1290 layered intrusions and the Keivitsa-Satovaara layered complex, northern Finland[END_REF]Törmänen et al., 473 2016). 474 475

Pt-Os isotope systematics 476

The high-precision Os isotopic data for the chromite separates from the Jeesiörova 477 komatiite samples 12D-PPR and 13 EJH are presented in Table 4 andplotted

W isotopic compositions and abundances 484

The W abundances and isotopic compositions of the Jeesiörova and Kevitsa komatiites are 485 listed in Tables 2 and5 The two Kevitsa komatiite samples and two replicates analyzed in this study for W 498 isotopic composition, are characterized by 182 W/ 184 W ratios that are indistinguishable from 499 182 W/ 184 W measured in the laboratory reference material, with an average  182 W = +1.5±3.3 500 (2SE, n = 4), where  182 W is the parts per million deviation of 182 W/ 184 W of a given sample 501 from that of the reference material, which, by definition, is equal to zero (Table 5). 502 503

Sm-Nd and Lu-Hf isotopic compositions and abundances 504

The 147 Sm-143 Nd isotopic and concentration data for the whole-rock samples and 505 clinopyroxene separates of the Jeesiörova and Kevitsa komatiites are listed in Table 6 and506 plotted in Fig. 9. Regression of the combined Sm-Nd isotopic data (Fig. 9a) for 13 samples 507 yields a Model 1 ISOPLOT internal isochron age of 2046±22 Ma (MSWD = 2.4), which is 508 identical, within the uncertainty, to the Re-Os isochron age of 2049±13 Ma. The initial  143 Nd 509 derived from the isochron is +3.8±0.9. The individual samples have very uniform initial Nd 510 isotopic compositions, with initial  143 Nd ranging between +3.5 and +3.9. A more precise 511 initial  143 Nd = +3.7±0.3 (2SD, n = 13) is obtained by averaging the initial 143 Nd/ 144 Nd ratios 512 calculated for each sample using the measured 147 Sm/ 144 Nd and 143 Nd/ 144 Nd ratios. Our results 513 are consistent with those obtained earlier for the most LREE-depleted Jeesiörova komatiites 514 [START_REF] Hanski | The Paleoproterozoic komatiite-1163 picrite association of Finnish Lapland[END_REF]. 515

The 176 Lu-176 Hf data (Fig. 9b) for the Jeesiörova whole-rock komatiite sample 13-EJH 516 and clinopyroxene separate define a regression line in 176 Lu/ 177 Hf-176 Hf/ 177 Hf space with a 517 slope corresponding to an ISOPLOT Model 1 age of 2072±37 Ma, which also overlaps, 518 within uncertainties, the Re-Os emplacement age of the Jeesiörova and Kevitsa komatiites. 519

The initial  176 Hf value derived from the regression is +8.7±0.7. The average initial  176 Hf 520 value calculated for the two samples analyzed (+8.7±0.3, 2SD, n = 2) is identical to the one 521 derived from the isochron. 522 Plotted in Fig. 9c are the Nd and Hf isotopic data for the Jeesiörova komatiites, together 523 with the Nd-Hf terrestrial array from Blichert-Toft and [START_REF] Blichert-Toft | Depleted mantle sources through time: Evidence from 1075 Lu-Hf and Sm-Nd isotope systematics of Archean komatiites[END_REF]. In contrast to other 524 late Archean and post-Archean komatiite systems (e.g., Kostomuksha: Blichert-Toft and 525 [START_REF] Blichert-Toft | Depleted mantle sources through time: Evidence from 1075 Lu-Hf and Sm-Nd isotope systematics of Archean komatiites[END_REF]Abitibi: Puchtel et al., 2018;Vetreny: Puchtel et al., 2016a;Gorgona: 526 Thompson et al., 2003), the calculated initial  143 Nd and  176 Hf values of the Jeesiörova 527 komatiites (Fig. 9c) do not plot on the terrestrial array of Blichert-Toft and [START_REF] Blichert-Toft | Depleted mantle sources through time: Evidence from 1075 Lu-Hf and Sm-Nd isotope systematics of Archean komatiites[END_REF]. 528 529 6. Discussion 530

Jeesiörova-Kevitsa emplaced komatiite magma composition 531

Nicklas et al. ( 2019) reported major and trace element abundance data for olivine 532 phenocrysts and whole-rock samples of the Kevitsa komatiites. Based on the Mg-Fe olivine-533 komatiite melt equilibria, these authors estimated that the parental magma to the Kevitsa 534 komatiites contained 27.4±2.4 wt.% MgO. By comparison for this study, in order to estimate 535 the parental magma composition, we used an approach different from that of Nicklas et al. 536 (2019). We performed mathematical incremental step-wise addition of olivine to a melt 537 composition represented by sample KD-14 (MgO = 16.3 wt.%) until the melt was in 538 equilibrium with olivine Fo 94.1 , the most MgO-rich olivine population found in the Kevitsa 539 komatiites (Nicklas et al., 2019, and this study). To select appropriate olivine-melt Fe-Mg 540 distribution coefficients (K D ), we utilized the results of the isothermal crystallization 541 experiments of komatiitic liquids carried out by [START_REF] Sossi | Liquidus temperatures of komatiites and the effect of 1403 cooling rate on element partitioning between olivine and komatiitic melt[END_REF]. Using the empirical 542 K D defined as (Fe 2+ /Mg) ol /(Fe/Mg) melt and the T-MgO-K D -logfO 2 relationships from their 543 study, we adjusted the K D values at each step of olivine addition. Because K D defined in this 544 way using bulk Fe in the melt is dependent on the oxygen fugacity, two different redox 545 conditions were considered, FMQ = +0.9 and FMQ = 0.0. The former choice is based on 546 the observed V partitioning between olivine and melt in the Kevitsa komatiites (Nicklas et al., 547 2019) and the latter on the composition of chrome spinel (Fe 3+ /(Fe 3+ +Al+Cr) = 0.06). The K D 548 values were varied between 0.28-0.26 and 0.29-0.27 in these two cases, respectively, and the 549 required additions of olivine were 34 and 40%. 550

The two resulting major element compositions of the emplaced komatiite melt are similar, 551 with the most important difference being the MgO content. The more oxidized conditions 552 yielded an MgO content of 25.2 wt.%, while the FMQ buffer conditions gave an MgO content 553 of 26.3 wt.%. These estimates are identical, within the uncertainties, to those obtained by 554 [START_REF] Nicklas | Secular Mantle Oxidation across the Archean-Proterozoic 1310 Boundary: Evidence from V partitioning in Komatiites and Picrites[END_REF] using an independent technique, and corroborate the komatiitic 555 composition of the emplaced Kevitsa komatiite magma. The results indicate that the parental 556 magmas to the Kevitsa, and, therefore, also Jeesiörova, komatiites approached those of their 557 late Archean counterparts from the Abitibi and Belingwe greenstone belts (Puchtel et al., 558 2009b) in terms of MgO content, and, therefore, their liquidus temperatures. 559 560

Lithophile trace element and isotope systematics of the komatiites 561

The Jeesiörova and Kevitsa komatiites have low (La/Sm) N (0.347±0.022 and 0.264±0.002, 562 respectively), and high positive initial Nd and Hf (+3.7±0.3 and +8.7±0.3, respectively), 563 indicating derivation from a mantle source that was strongly depleted in highly incompatible, 564 lithophile trace elements. These komatiites are ~2.5 times more strongly depleted in the light 565 REE than, e.g., modern N-MORB (average (La/Sm) N = 0.654 [START_REF] Hofmann | Chemical differentiation of the Earth: The relationship between 1170 mantle, continental crust and oceanic crust[END_REF]). Yet, the 566 lavas are also characterized by enrichments in Th, U, and W relative to Nb, an element with 567 similar incompatibility during melting of spinel peridotite ((Nb/Th) N = 0.583±0.011), and a 568 negative Nb anomaly (Nb/Nb* = 0.582±0.010, where Nb/Nb* = Nb N /√(Th N *La N )). This is 569 most likely indicative of minor, yet significant crustal contamination. 570

Upper crustal rocks are generally characterized by low (Nb/Th) N and (Nb/La) N ratios and, 571 as a result, display pronounced negative Nb anomalies relative to the elements with similar 572 incompatibility during mantle melting (Th and La) on BSE-normalized trace element 573 diagrams, i.e., their Nb/Nb* <<1.0 [START_REF] Rudnick | Nature and composition of the continental crust: a 1385 lower crustal perspective[END_REF][START_REF] Rudnick | Composition of the Continental Crust[END_REF]. 574 By contrast, primary uncontaminated komatiitic magmas commonly do not show a Nb 575 anomaly, i.e., their Nb/Nb* = 1.00 [START_REF] Jochum | Nb-Th-La in komatiites and basalts: 1200 constraints on komatiite petrogenesis and mantle evolution[END_REF][START_REF] Puchtel | Oceanic plateau model for continental crustal growth in the Archaean: A case 1326 study from the Kostomuksha greenstone belt, NW Baltic Shield[END_REF]. Jochum et al. 576 (1991) noted that the elements most sensitive to crustal contamination in komatiite magmas 577 include Th, U, and light REE and, to a lesser extent, Nb. As a result, the Nb/Nb* ratio can be 578 utilized as a measure of the degree of crustal contamination and, when applied in combination 579 with the lithophile element isotopic systems, to discriminate between contaminated and 580 uncontaminated lavas. Based on these observations, it is likely that the original Jeesiörova-581 Kevitsa komatiite magma was contaminated with upper crustal rocks of the Fennoscandian 582 Shield en route to the surface. 583

In order to test this hypothesis and calculate the major and trace element and Nd, Hf, and 584 W isotopic compositions of the original Jeesiörova-Kevitsa komatiite magma, we performed 585 model mixing calculations, the results of which are presented in Table 8. For the crustal 586 contaminant endmember, we assumed an average composition of the Vodla Block tonalitic 587 gneisses that are the dominant upper crustal rock type within the adjacent Karelian granite-588 greenstone terrane (Puchtel et al., 2016b). These Karelian tonalites are similar in age and 589 major-and trace element composition to their counterparts from Finnish Lapland (e.g., 590 [START_REF] Huhma | The age of the 1182 Archean greenstone belts in Finland[END_REF], but, unlike the latter, have been better characterized in terms of the 591 isotopic systematics of interest (Table 8). This average tonalite composition from Puchtel et 592 al. (2016b) is henceforth referred to as the Fennoscandian Tonalite Average (FTA). In order 593 to calculate the degree of contamination and the trace element abundances and Nd and Hf 594 isotopic compositions of the original, uncontaminated Jeesiörova-Kevitsa komatiite magma, 595 we assumed that it had a Nb/Nb* = 1.00. Calculations show that to produce the emplaced 596 Jeesiörova-Kevitsa komatiite magma with Nb/Nb* = 0.582, the original komatiite magma had 597 to be mixed with ~1% of the average FTA (Nb/Nb* = 0.118). 598

Using 1% contamination and the Nd and Hf abundances and initial Nd and Hf isotopic 599 compositions of the FTA calculated at the time of the Kevitsa komatiite magma emplacement, 600 the original Kevitsa komatiite magma and, thus, its source is calculated to have had initial 601 Nd and Hf values of +4.9 and +10.2, respectively. 602

On the Nd versus Hf evolution diagram in Fig. 9c, the calculated datum for the 603 Jeesiörova-Kevitsa komatiite source plots within uncertainty of the terrestrial evolution curve 604 of Blichert-Toft and [START_REF] Blichert-Toft | Depleted mantle sources through time: Evidence from 1075 Lu-Hf and Sm-Nd isotope systematics of Archean komatiites[END_REF]. This indicates coupled, or congruent, behavior of the two 605 isotope systems in the source of the original komatiite magma. It appears, therefore, that the 606 Nd-Hf isotope decoupling observed in the Jeesiörova-Kevitsa emplaced komatiite magma was 607 caused by contamination with the older tonalitic crust that had isotopically evolved away 608 from the terrestrial evolution curve by the time it was mixed with the original komatiite 609 magma (Table 8). The datum for the Jeesiörova-Kevitsa komatiite source also plots near the 610 Depleted MORB Mantle (DMM) value of [START_REF] Goldstein | Nd and Sr isotopic systematics of river water 1148 suspended material: implications for crustal evolution[END_REF] at 2049 Ma (Fig. 611 9c), indicating derivation from a mantle source that was already as long-term depleted in 612 incompatible lithophile trace elements as DMM at 2.05 Ga. 613 614

Osmium isotope -HSE systematics of the Jeesiörova-Kevitsa komatiite source 615

Although the Os isotopic compositions and absolute and relative HSE abundances in the 616 mantle sources of early and late Archean komatiites have been the subject of numerous 617 studies over nearly four decades (e.g., [START_REF] Brügmann | Noble metal 1095 abundances in komatiite suites from Alexo, Ontario, and Gorgona Island, Colombia[END_REF][START_REF] Walker | Comparative Re-Os, Sm-Nd and Rb-Sr 1449 isotope and trace element systematics for Archean komatiite flows from Munro Township, 1450 Abitibi belt, Ontario[END_REF]Foster et 618 al., 1996;[START_REF] Gangopadhyay | Re-Os systematics of the ca. 2.7 Ga komatiites 1135 from Alexo, Ontario, Canada[END_REF]Puchtel et al., 2009a;2014, 2016b;Maier et al., 619 2009), studies of post-Archean komatiites have been rare, in part because the number of 620 Proterozoic and younger komatiites is far more limited. 621

In order to characterize the HSE present in the Paleoproterozoic mantle source of the 622 Jeesiörova-Kevitsa komatiites, the effects of crustal contamination must first be evaluated. 623

We again use the average FTA composition from Puchtel et al. (2016b) to assess the effects 624 of 1% contamination of the original Jeesiörova-Kevitsa komatiite magma on its Os isotopic 625 composition and HSE abundances. Because of the very low concentrations of these elements 626 in the FTA, the effects are negligible. Contamination of this magnitude would have increased 627

 186 Os and  187 Os values by only 0.004 (based on the calculated 186 Os/ 188 Os = 0.1198767 at 628 2.05 Ga for the FTA assuming its derivation from a BSE-type mafic source at 3.2 Ga, per 629 Puchtel et al., 2016b) and 0.01 (based on the calculated 187 Os/ 188 Os = 1.2091 at 2.05 Ga for 630 the FTA assuming its derivation from a BSE-type mafic source at 3.2 Ga, per Puchtel et al., 631 2016b) units, respectively, which are well within the uncertainties of the current estimates for 632 these values. Similarly, effects on absolute HSE abundances would have been <<1%. 633

Estimation of the HSE abundances in the mantle source of the Jeesiörova-Kevitsa 634 komatiites is achieved by the bootstrap method of Puchtel et al. (2004b) that sequentially 635 combines Os isotopic and HSE concentration data. First, we used the initial 187 Os/ 188 Os and 636

186 Os/ 188 Os ratios calculated for the Jeesiörova-Kevitsa komatiites to estimate the time-637 integrated evolution of Re/Os and Pt/Os in their mantle source. In order to place minimum 638 constraints on these long-term source ratios, we calculated the parent/daughter elemental 639 ratios necessary to arrive at the Os isotopic composition at 2049 Ma by assuming generation 640 of this mantle domain soon after Solar System formation. It is estimated that this source 641 evolved from an early Solar System 187 Os/ 188 Os = 0.09517 at 4567 Ma [START_REF] Archer | Highly siderophile elements 1041 and 187 Re-187 Os isotopic systematics of the Allende meteorite: Evidence for primary 1042 nebular processes and late-stage alteration[END_REF] to 642 the initial 187 Os/ 188 Os = 0.11285±23 at 2049 Ma with 187 Re/ 188 Os = 0.397±5. This ratio is well 643 within the range of bulk chondritic meteorites. Using the early Solar System 186 Os/ 188 Os = 644 0.1198269 at 4567 Ma [START_REF] Brandon | Platinum-osmium isotope evolution of 1090 the Earth's mantle: Constraints from chondrites and Os-rich alloys[END_REF] requires a source with a 190 Pt/ 188 Os = 0.00269±8 645 to have evolved to its 186 Os/ 188 Os = 0.1198369±3 at 2049 Ma. In contrast to Re/Os, the 646 required minimum 190 Pt/ 188 Os ratio is ~50% higher than the average 190 Pt/ 188 Os = 0.00180±17 647 (2SD) in bulk chondritic meteorites (as compiled from the data of [START_REF] Horan | Highly 1178 siderophile elements in chondrites[END_REF], 648 [START_REF] Brandon | Re-Os isotopic systematics 1087 and platinum group element composition of the Tagish Lake carbonaceous chondrite[END_REF][START_REF] Brandon | Platinum-osmium isotope evolution of 1090 the Earth's mantle: Constraints from chondrites and Os-rich alloys[END_REF], [START_REF] Fischer-Gödde | Rhodium, gold and other highly 1112 siderophile element abundances in chondritic meteorites[END_REF][START_REF] Van Acken | High-precision osmium isotopes in 1439 enstatite and Rumuruti chondrites[END_REF]. 649

The calculated initial  187 Os of the Jeesiörova-Kevitsa komatiite source is also within the 650 range of those for the majority of Archean and Paleoproterozoic komatiite sources, which 651 were characterized by initial  187 Os values between -0.1 and +1.3, also well within the range 652 of 95% of chondritic meteorites (Fig. 10a). The 186 Os/ 188 Os isotope data available for 653 Archean and Paleoproterozoic komatiite sources indicate that mantle sources of the 2.7 Ga 654 Abitibi (Canada), 2.7 Ga Belingwe (South Africa), and 2.4 Ga Vetreny (Fennoscandia) 655 komatiite systems evolved with time-integrated Pt/Os within the chondritic range (Fig. 10b), 656 whereas the coupled high initial 186,187 Os/ 188 Os of the 2.8 Ga Kostomuksha komatiite system 657 (Fennoscandia) require long-term suprachondritic Pt/Os and Re/Os (Puchtel et al., 2005). By 658 contrast, the 3.5 Ga Komati and 3.3 Ga Weltevreden komatiite systems (South Africa) 659 evolved with non-chondritic Pt/Os, but chondritic Re/Os ratios [START_REF] Puchtel | Insights into 1358 Early Earth from the Pt-Re-Os isotope and Highly Siderophile Element abundance 1359 systematics of Barberton komatiites[END_REF], 660 displaying decoupling of the Re-Os and Pt-Os isotopic systems. The calculated 661 suprachondritic initial  186 Os = +29±2 in the Jeesiörova-Kevitsa komatiite source obtained in 662 this study is similar to  186 Os = +22±7 in the source of the Weltevreden komatiite system 663 [START_REF] Puchtel | Insights into 1358 Early Earth from the Pt-Re-Os isotope and Highly Siderophile Element abundance 1359 systematics of Barberton komatiites[END_REF]. 664

Because of the availability of high-precision Os isotopic data and immobile behavior of 665 HSE in the Kevitsa komatiites during post-magmatic alteration, absolute abundances of 666 incompatible Re, Pt, and Pd and compatible Os in the Jeesiörova-Kevitsa komatiite source 667 can be estimated using the bootstrap protocol of Puchtel et al. (2004b). 668

As was shown in a number of previous studies (e.g., [START_REF] Alard | Non-chondritic 1028 distribution of the highly siderophile elements in mantle sulfides[END_REF]Lorand and Alard, 669 2001;[START_REF] Luguet | Residual 1247 platinum-group minerals from highly depleted harzburgites of the Lherz massif (France) 1248 and their role in HSE fractionation of the mantle[END_REF], ≥90% of the HSE budget of the mantle resides in two major types 670 of sulfides. The high-temperature Os-Ir-Ru-rich Fe-Ni monosulfide solid solution (mss) 671 usually occurs as inclusions in liquidus phases, primarily olivine, whereas lower-temperature 672 Cu-Ni sulfides occupy interstitial space between the liquidus phases. During partial melting of 673 mantle peridotite, Cu-Ni sulfides enter the melt, whereas mss remains trapped in the melting 674 residue. It takes ~18-25% partial melting [START_REF] Barnes | The origin of the fractionation of platinum-1057 group elements in terrestrial magmas[END_REF][START_REF] Keays | The role of komatiitic and picritic magmatism and S-saturation in the 1206 formation of ore deposits[END_REF]Luguet et al., 675 2007;[START_REF] Mungall | Ore deposits of the platinum-group elements[END_REF][START_REF] Naldrett | Secular Variation of Magmatic Sulfide Deposits and Their Source 1299 Magmas[END_REF][START_REF] Fonseca | Solubility of Os and Ir in sulfide melt: Implications for Re/Os fractionation during mantle 1125 melting[END_REF]2012;Kiseeva et al., 676 2017) for all the low-temperature Cu-Ni sulfides to get exhausted in the mantle. As the degree 677 of melting continues to increase, the magma becomes sulfide-undersaturated. It has also 678 recently been shown that decrease in fS 2 with increase in the degree of melting triggers 679 exsolution of Os-Ir alloys from the refractory mss in the residue [START_REF] Fonseca | Solubility of Os and Ir in sulfide melt: Implications for Re/Os fractionation during mantle 1125 melting[END_REF]2012). 680

All low-degree (basalts) and the majority of higher-degree (picrites and komatiites) partial 681 melts are characterized by compatible behavior of Os and Ir during magmatic differentiation, 682 indicating that their parental magmas remained saturated in Os-Ir alloys (Puchtel et al., 683 2004b;[START_REF] Barnes | Iridium, ruthenium and rhodium in komatiites: 1061 Evidence for iridium alloy saturation[END_REF]. However, some lavas, such as the 2.8 Ga Kostomuksha 684 and the 3.55 Ga Schapenburg komatiites, exhibited an incompatible behavior of Os and Ir 685 during magma differentiation, likely indicating near-complete exhaustion of Os-Ir alloys in 686 the mantle sources of these komatiites (Puchtel and Humayun, 2005;Puchtel et al., 2009a). 687

Apparently, the sulfide saturation status of komatiite magmas is a crucial parameter that 688 plays a key role in determining whether their HSE abundances accurately reflect those in their 689 mantle sources. In the following discussion, we evaluate the sulfide saturation status of the 690 Jeesiörova-Kevitsa komatiite magma during both mantle melting and after emplacement. 691 Sulfide saturation of a mafic magma is determined by several key parameters, the most 692 important of which are pressure (P), temperature (T), and FeO content. In their recent study, 693 [START_REF] Steenstra | An experimental assessment of the potential of sulfide saturation of the source 1407 regions of eucrites and angrites: Implications for asteroidal models of core formation, late 1408 accretion and volatile element depletions[END_REF] combined the results of their own work with previously published 694 experimental data of [START_REF] Mavrogenes | The relative effects of pressure, temperature 1261 and oxygen fugacity on the solubility of sulfide in mafic magmas[END_REF], [START_REF] Li | The effect of pressure, temperature, oxygen fugacity and 1238 composition on partitioning of nickel and cobalt between liquid Fe-Ni-S alloy and liquid 1239 silicate: implications for the earth's core formation[END_REF], Holzheid and 695 Grove (2002), O'Neill and [START_REF] O'neill | The Sulfide Capacity and the Sulfur Content at 1313 Sulfide Saturation of Silicate Melts at 1400°C and 1 bar[END_REF], [START_REF] Jugo | An Experimental Study of the Sulfur 1203 Content in Basaltic Melts Saturated with Immiscible Sulfide or Sulfate Liquids at 1300°C 1204 and 1.0 GPa[END_REF], [START_REF] Liu | Sulfur concentration at sulfide saturation 1242 (SCSS) in magmatic silicate melts[END_REF], Brenan 696 (2008), Kiseeva andWood (2013, 2015), [START_REF] Ding | Sulfur concentration of martian basalts at sulfide 1106 saturation at high pressures and temperatures -Implications for deep sulfur cycle on Mars[END_REF][START_REF] Ding | New high pressure experiments on sulfide saturation 1109 of high-FeO*-basalts with variable TiO 2 contents -Implications for the sulfur inventory 1110 of the lunar interior[END_REF], [START_REF] Fortin | The effect of water on 1127 the sulfur concentration at sulfide saturation (SCSS) in natural melts[END_REF], Wood 697 and Kiseeva (2015), [START_REF] Smythe | The S content of silicate melts at sulfide 1400 saturation: New experiments and a model incorporating the effects of sulfide composition[END_REF], and [START_REF] Steenstra | Evidence for a sulfur-undersaturated 1411 lunar interior from the solubility of sulfur in lunar melts and sulfide-silicate partitioning of 1412 siderophile elements[END_REF] to produce a refined 698 parameterization model for the sulfur concentration at sulfide saturation (SCSS) as a function 699 of P, T, and major and trace element abundances in silicate liquids (Eq. 4 in Steenstra et al., 700 2020). 701

In order to calculate the P-T parameters of the Jeesiörova-Kevitsa komatiite magma 702 generation, we used the mantle melting parameterization model of McKenzie and Bickle 703 (1988). These calculations yield the following parameters: potential mantle temperature of the 704 Jeesiörova-Kevitsa komatiite mantle source T pot = 1763°C, liquidus temperature upon 705 emplacement T liq = 1544°C, depth of melting initiation 294 km, which translates into a 706 pressure of 9.5 GPa. Based on these parameters, the temperatures and pressures at which 707 melting terminated for the Jeesiörova-Kevitsa komatiite magma was calculated using the 708 model of [START_REF] Herzberg | Thermal history of the Earth and its 1167 petrological expression[END_REF] to be 1730°C and 5.0 GPa. Using these parameters, the 709 chemical composition of the original Jeesiörova-Kevitsa komatiite magma from Table 8, and710 Eq. 4 of [START_REF] Steenstra | An experimental assessment of the potential of sulfide saturation of the source 1407 regions of eucrites and angrites: Implications for asteroidal models of core formation, late 1408 accretion and volatile element depletions[END_REF], the SCSS of the Jeesiörova-Kevitsa komatiite magma upon 711 separation from the melting residue and ensuing emplacement are calculated to be 2270 and 712 5200 ppm, respectively. 713

In order to estimate the S content in the Jeesiörova-Kevitsa komatiite magma and, thus, its 714 sulfide saturation status, knowledge of the sulfur content of its mantle source, the S partition 715 coefficient, and the degree of melting is required. The average S content of modern DMM is 716 estimated to be 206±25 ppm [START_REF] Sun | 1416 Sulfur abundance and heterogeneity in the MORB mantle estimated by copper 1417 partitioning and sulfur solubility modelling[END_REF]. Since the Jeesiörova-Kevitsa komatiite source 717 was as strongly melt-depleted as contemporary DMM, we assume it also had a similar S 718 content. 719

To estimate the degree of partial melting that produced the original Jeesiörova-Kevitsa 720 komatiite magma, we used a batch melting model and the calculated abundances of moderately incompatible lithophile elements (e.g., Al, Ti, Gd) in this magma and its garnet-722 free mantle source of spinel lherzolite composition, as indicated by their Al-undepleted nature 723 and the depth of melting initiation of ~290 km. These calculations yield 49% partial melting. 724

Assuming that sulfur was highly incompatible during komatiite melting, with the degree of 725 incompatibility similar to that of La, the original Jeesiörova-Kevitsa komatiite magma is 726 calculated to have contained 420 ppm S. This S content is, thus, ~5 times lower than the 727 SCSS of the Jeesiörova-Kevitsa komatiite magma upon separation from the melting residue 728 and ~12 times lower than the SCSS of the Jeesiörova-Kevitsa komatiite magma upon 729 emplacement, indicating that it was strongly sulfur-undersaturated upon separating from the 730 source and remained as such all the way to the surface. 731

The sulfur-undersaturated nature of the Jeesiörova-Kevitsa komatiite magma upon 732 emplacement is also attested to by the incompatible behavior of the highly chalcophile 733 element Cu during magma differentiation after emplacement. The Cu abundances in the 734 Kevitsa komatiites strongly correlate with the MgO content and plot on an olivine control line 735 in the MgO versus Cu diagram (Fig. 4a), indicating that sulfide liquid was not a fractionating 736 phase within the compositional range represented by the samples of this study. 737

To estimate the abundances of incompatible Pt and Pd in the Jeesiörova-Kevitsa komatiite 738 source, we extrapolated the differentiation trends for each of these elements to a mantle MgO 739 = 38 wt.% using ISOPLOT. The concentrations of Pt and Pd in the source were, thus, 740 calculated to be 9.54±0.22 and 8.13±0.19 ppb, respectively, and the total Pt and Pd 741 abundances to be 17.7±0.8 ppb, or 120±5% of those in the average estimates for modern BSE 742 of 14.7±2.0 ppb [START_REF] Becker | Highly 1063 siderophile element composition of the Earth's primitive upper mantle: Constraints from 1064 new data on peridotite massifs and xenoliths[END_REF]. 743

The concentration of compatible Os in the source was estimated to be 3.4±0.1 ppb using 744 the calculated Pt content in the source and the long-term 190 Pt/ 188 Os derived from the initial 745 186 Os/ 188 Os. Further, using the calculated Os concentration and the estimated long-term 746 187 Re/ 188 Os derived from the initial 187 Os/ 188 Os, the Re concentration in the source is 747 calculated to be 0.285±0.008 ppb. Unlike the Pt and Pd concentrations, therefore, the 748 calculated Os and Re concentrations in the source of the Jeesiörova-Kevitsa komatiites are, 749 respectively, 13% and 19% lower than the average estimates for BSE. 750

The presence of residual Os/Ir alloys in the melting residue, as evidenced by the 751 compatible behavior of Os and Ir during the Jeesiörova-Kevitsa magma differentiation, makes 752 it possible for the Os/Ir ratio to fractionate between the source and the komatiitic magma. 753 Furthermore, as discussed earlier, chromite fractionation likely had a strong effect on Ru 754 concentrations in the Jeesiörova-Kevitsa komatiite magma. Based on these observations, we 755 were unable to constrain the abundances of Ir and Ru in the source with any meaningful 756 degree of accuracy. 757

The calculated total Pt and Pd abundances in the Jeesiörova-Kevitsa komatiite source are 758 also plotted as a function of age in Fig. 11 and compared with the data for other Archean and 759 Paleoproterozoic komatiite systems studied to date and the BSE estimate. The calculated total 760 Pt and Pd abundances in the source of the Jeesiörova-Kevitsa komatiites are significantly 761 higher than in any other komatiite system examined so far. The significance of this 762 observation will be discussed in the following sections. 763

The siderophile and lithophile element isotope conundrum 764

There is an apparent decoupling between the radiogenic lithophile trace element isotope 765 systematics (e.g., Sm-Nd and Lu-Hf) and Re-Os isotope systematics in komatiite mantle 766 sources worldwide. Many komatiite systems, such as those from the Abitibi (Canada) and 767

Belingwe (South Africa) greenstone belts (Puchtel et al., 2009b;[START_REF] Puchtel | 182 W and HSE 1368 constraints from 2.7 Ga komatiites on the heterogeneous nature of the Archean mantle[END_REF], the 768 Weltevreden and Komati (South Africa) Formations [START_REF] Puchtel | Insights into 1358 Early Earth from the Pt-Re-Os isotope and Highly Siderophile Element abundance 1359 systematics of Barberton komatiites[END_REF], and the Sumozero-769 Kenozero (Fennoscandia) greenstone belt [START_REF] Puchtel | Os-Pb-Nd isotope and highly siderophile 1344 and lithophile trace element systematics of komatiitic rocks from the Volotsk suite, SE 1345 Baltic Shield[END_REF], are characterized by strong, 770 long-term depletions in incompatible lithophile trace elements, e.g., Nd versus Sm, but 771 chondritic Re/Os. By contrast, certain other komatiite systems, such as the Kostomuksha in 772 Fennoscandia (Puchtel et al., 2005) and the Schapenburg in South Africa (Puchtel et al., 773 2016a), while being long-term depleted in light REE, evolved with long-term suprachondritic 774 Re/Os ratios. In this study, we address this apparent conundrum using the Sm-Nd, Lu-Hf, and 775

Re-Os isotope and lithophile trace element abundance data for the Jeesiörova-Kevitsa 776 komatiite system. 777

In Section 6.2, we calculated that the mantle source of the Jeesiörova-Kevitsa komatiites 778 had the initial  143 Nd = +4.9 and  176 Hf = +10.2. This source was similar to DMM at 2.05 Ga 779 [START_REF] Goldstein | Nd and Sr isotopic systematics of river water 1148 suspended material: implications for crustal evolution[END_REF] Jeesiörova-Kevitsa komatiite source evolved within the range of chondritic Re/Os and 787 suprachondritic Pt/Os. 788

Tungsten isotopic composition of the Jeesiörova-Kevitsa komatiite source 789

The Paleoproterozoic mantle was recently shown to have been characterized by an 790 oxidation state similar to that of the modern mantle [START_REF] Nicklas | Secular Mantle Oxidation across the Archean-Proterozoic 1310 Boundary: Evidence from V partitioning in Komatiites and Picrites[END_REF]. As a result, during 791 melting of the Paleoproterozoic mantle, W likely was characterized by a degree of 792 incompatibility similar to that of Th and U [START_REF] Newsom | The depletion of tungsten in the bulk silicate Earth: Constraints on core formation[END_REF]Arevalo and McDonough, 793 2008;[START_REF] König | The Earth's tungsten budget during mantle melting and crust formation[END_REF]. Therefore, the BSE-normalized W abundance in the emplaced 794 Jeesiörova-Kevitsa komatiite magma is expected to be similar to those of Th and U. This is 795 the case, with an average calculated W/W* = 1.0. This also implies that W, similarly to Th 796 and U, was characterized by immobile behavior during the limited postmagmatic alteration of 797 the Jeesiörova-Kevitsa komatiites. Further, similarly to the other lithophile trace elements, 798 such as Nd and Hf, upper crustal rocks have significantly higher W abundances than 799 komatiites; they also likely had a different W isotopic composition compared to komatiites, 800 and, therefore, addition of even a small amount of upper crustal material can significantly 801 change both the W isotopic composition and W abundances of a komatiite magma. 802

In this study, we used the measured W isotopic composition and W abundance of the 803 emplaced Jeesiörova-Kevitsa komatiite magma to calculate those of the original, non-804 contaminated komatiite magma. In order to remove the effect of the 1% crustal contamination 805 estimated on the basis of the lithophile trace element systematics, we used the W isotopic 806 composition and W abundances of the average FTA from Puchtel et al. (2016b). The 807 endmember parameters and the results of the mixing calculations are given in Table 8. The 808 calculations suggest that the original Jeesiörova-Kevitsa komatiite magma contained ~5.8 ppb 809 W, implying that ~50% of the total W budget in the emplaced Jeesiörova-Kevitsa komatiite 810 magma was derived from the crustal contaminant. 811

The average  182 W value of the FTA reported by Puchtel et al. (2016b) was +12.6±4.5, so 812 contamination of the original komatiite magma with W-rich crust would have led to an 813 increase in the  182 W value of the original magma. If the assumptions regarding the original 814 uncontaminated komatiite magma concentration of W, and concentration and 182 W isotopic 815 composition of the crustal contaminant are accurate, the W isotopic composition of the 816 original Jeesiörova-Kevitsa komatiite magma and, by extension, of its mantle source, after 817 correction for crustal contamination, is calculated to be  182 W = -10±5. Here, the uncertainty 818 solely reflects propagated measurement uncertainties for the komatiite and the FTA. Because 819 it is a projected original melt composition, this  182 W must be considered tentative. 820 Nevertheless, negative  182 W values of this magnitude are common in modern ocean island 821 basalt systems [START_REF] Rizo | 182 W evidence for core-mantle 1376 interaction in the source of mantle plumes[END_REF][START_REF] Mundl-Petermeier | Temporal evolution of primordial tungsten-182 and 3 He/ 4 He 1283 signatures in the Iceland mantle plume[END_REF]2020) and have also been 822 observed in the 3.6 Ga Schapenburg komatiites (Puchtel et al., 2016a), South Africa, and 823 some spatially associated, Archean glacial diamictites [START_REF] Mundl | Tungsten-1279 182 in the upper continental crust: Evidence from glacial diamictites[END_REF]. By contrast, the 824 majority of evolved Archean crustal rocks similar to those envisioned for the FTA 825 contaminant, are characterized by positive 182 W anomalies (e.g., [START_REF] Touboul | New Insights into the 1429 Hadean Mantle Revealed by 182 W and Highly Siderophile Element Abundances of 1430 Supracrustal Rocks from the Nuvvuagittuq Greenstone Belt, Quebec, Canada[END_REF]826 Willbold et al., 2015;Rizo et al., 2016b). 827

Reconciling isotopic characteristics of the Jeesiörova-Kevitsa komatiites 828

The combined chemical and isotopic data for the Jeesiörova-Kevitsa komatiites are 829 complex. The Nd and Hf isotopic data are consistent with derivation from a mantle domain 830 that experienced at least one melting event prior to the one that produced the komatiitic 831 magma. The Nd-Hf isotope characteristics of this mantle domain are similar to projections of 832 the modern DMM to 2049 Ma. The 187 Os isotopic data indicate derivation from a broadly 833 chondritic mantle source that was not strongly long-term melt depleted. Because the long-834 term effects of low degrees of partial melting on Re/Os, compared to Sm/Nd and Lu/Hf, are 835 relatively small (Puchtel et al., 2004b;2009b), the chondritic 187 Os/ 188 Os of the mantle source 836 can be considered at least qualitatively consistent with the Nd and Hf isotopic data. By 837 contrast, the fractionated Pt/Os of the mantle source, as revealed by the suprachondritic initial 838

186 Os/ 188 Os, and the correspondingly fractionated HSE pattern for the source, requires an 839 event that could have strongly modified HSE relative abundances early in Earth history. 840 Finally, if the negative  182 W value for the Jeesiörova-Kevitsa komatiite system obtained 841 after correction for crustal contamination is accurate, it requires invoking one of a limited 842 number of possible processes to account for the projected anomalous composition for the Hf-843 W short-lived radiogenic isotope system. 844

To address these complexities, we consider two melting models. The parameters used in 845 both models are as follows. (1) The mantle domain that ultimately became the source of the 846 Jeesiörova-Kevitsa komatiites initially had a composition consistent with that of the BSE. The 847 chemical and isotopic parameters of this domain were compiled from Jacobsen and 848 Wasserburg (1980), [START_REF] Hamilton | Sm-Nd studies of 1153 Archaean metasediments and metavolcanics from West Greenland and their implications 1154 for the Earth's early history[END_REF]), Bouvier et al. (2008[START_REF] Hofmann | Chemical differentiation of the Earth: The relationship between 1170 mantle, continental crust and oceanic crust[END_REF], Shirey 849 and Walker (1998), [START_REF] Brandon | Platinum-osmium isotope evolution of 1090 the Earth's mantle: Constraints from chondrites and Os-rich alloys[END_REF][START_REF] Becker | Highly 1063 siderophile element composition of the Earth's primitive upper mantle: Constraints from 1064 new data on peridotite massifs and xenoliths[END_REF]. ( 2) This domain had a 850 modal mineral composition typical of spinel lherzolite, with 65% olivine, 24% orthopyroxene, 851 10% clinopyroxene, and 1% spinel. (3) The melting of the domain was modeled using the 852 batch melting equation of Shaw (1970). The bulk D Sm , D Nd , D Lu , and D Hf during melting were 853 calculated to be 0.040, 0.024, 0.090, and 0.037, respectively, based on the above source 854 modal composition and crystal-liquid D values for the respective mineral phases from Green 855 (1994). The bulk D Re and D Pt under the redox conditions of FMQ +0.90±0.33 determined by 856 [START_REF] Nicklas | Secular Mantle Oxidation across the Archean-Proterozoic 1310 Boundary: Evidence from V partitioning in Komatiites and Picrites[END_REF] were estimated to be 0.20 and 2.0, respectively (Rehkämper et al., 1999;857 Mallmann and O'Neill, 2007). The bulk D Os for the low-degree partial melting, consistent 858 with formation of a MORB-type magma, was adopted to be 50 (e.g., Gannoun et al., 2016 and 859 references therein). 860

In Model 1, we consider an early, global differentiation event. This model requires a 861 minimum degree of melting and a melt extraction that would lead to fractionation of Sm/Nd 862 and Lu/Hf in the BSE domain from the chondritic value to that required to bring the  143 Nd 863 and  176 Hf of the domain to the corrected values of +4.9 and +10.2, respectively, by the time 864 of komatiite formation at 2049 Ma. The effect of this melt extraction event on the Re-Os and 865

Pt-Os isotope systems is then considered. Model 2 considers the youngest-possible, maximum 866 degree of melting event that would fractionate the Sm/Nd and Lu/Hf ratios in the BSE domain 867 from the chondritic values to those calculated for the Jeesiörova-Kevitsa komatiite source, 868 and again bring the  143 Nd and  176 Hf values of this domain to +4.9 and +10.2, respectively, 869 by 2049 Ma. The effect of this melt extraction event on the Re-Os and Pt-Os isotope systems 870 is then also considered. 871

For Model 1, we assume that the global differentiation event occurred 100 Ma after Solar 872 System formation. This timing is consistent with the recent data on the early formation of the 873 Moon at ~4.51 Ga [START_REF] Barboni | Early formation of the Moon 4.51 billion years ago[END_REF]. The Jeesiörova-Kevitsa komatiite source would 874 have evolved from chondritic ratios of 143 Nd/ 144 Nd = 0.506807 and 176 Hf/ 177 Hf = 0.279863 at 875 4.467 Ga to the initial ratios of 143 Nd/ 144 Nd = 0.510236±8 and 176 Hf/ 177 Hf = 0.281762±18 at 876 2049 Ma with long-term ratios of 147 Sm/ 144 Nd = 0.2123±5 and 176 Lu/ 177 Hf = 0.0394±4. Such 877 modest change in 147 Sm/ 144 Nd and 176 Lu/ 177 Hf from the chondritic ratios of 0.1967 and 878 0.0336, respectively, would require only ~0.6% melting, followed by removal of the melt. At 879 the same time, this melt extraction event would have decreased the 187 Re/ 188 Os in the melting 880 residue from 0.4019 to 0.3912; the latter ratio is essentially identical to the calculated long-881 term 187 Re/ 188 Os = 0.397±8 determined for the Jeesiörova-Kevitsa komatiite source. However, 882 such low-degree melting event would not have had any measurable effect on the Pt/Os ratio 883 of the source. Therefore, Model 1 cannot explain the observed 186 Os/ 188 Os composition of the 884 Jeesiörova-Kevitsa komatiite system. 885

For Model 2, using the lithophile trace element data, we first calculate the 147 Sm/ 144 Nd 886 and 176 Lu/ 177 Hf ratios in the Jeesiörova-Kevitsa komatiite source to be 0.2912 and 0.0704, 887 respectively. Modeling indicates that 9% melting would have generated a residual mantle 888 domain with these Sm/Nd and Lu/Hf ratios. The latest time of the melting event required for 889 the mantle domain with such Sm/Nd and Lu/Hf ratios to have evolved to the initial  143 Nd = 890 +4.9 and  176 Hf = +10.2 by 2049 Ma is 2.45 Ga. Any event that occurred later than that would 891 not have afforded sufficient time to develop such radiogenic  143 Nd and  176 Hf in the mantle 892 domain. It is noteworthy that 2.45 Ga coincides with the ages of abundant mafic layered 893 intrusions [START_REF] Alapieti | Early Proterozoic layered intrusions in the northeastern part of the 1026 Fennoscandian Shield[END_REF][START_REF] Balashov | Isotope data on the age and 1051 genesis of layered basic-ultrabasic intrusions in the Kola Peninsula and northern Karelia, 1052 northeastern Baltic Shield[END_REF][START_REF] Amelin | U-Pb geochronology of layered 1034 mafic intrusions in the eastern Baltic Shield: implications for the timing and duration of 1035 Paleoproterozoic continental rifting[END_REF]Yang et al., 2016) 894 and komatiitic lavas [START_REF] Puchtel | Petrology and geochemistry of crustally contaminated 1321 komatiitic basalts from the Vetreny Belt, southeastern Baltic Shield: Evidence for an early 1322 Proterozoic mantle plume beneath rifted Archean continental lithosphere[END_REF]2016b) throughout the Fennoscandian Shield, 895 indicating that large volumes of mafic magma were indeed produced in this part of the planet 896 at that time. 897 Using this extent and time of melting, and applying a bulk D Re = 0.2, model calculations 898 indicate that the melting event that produced this light REE depletion would have reduced the 899 187 Re/ 188 Os ratio in the chondritic reservoir from 0.4019 to 0.2694 at 2.45 Ga. The resultant 900 melt-depleted mantle domain, if allowed to evolve from 2.45 Ga until the time of the 901 komatiite formation at 2049 Ma with this Re/Os ratio, would have developed a negative initial 902  187 Os = -0.8, which is close to the observed value and well within the range of chondritic 903 meteorites at 2049 Ma (Fig. 10). At the same time, this melting event would only have 904 reduced the 190 Pt/ 188 Os ratio in the mantle domain by 0.5%. The 400 Ma of growth from 2.45 905 to 2.05 Ga with this slightly reduced Pt/Os ratio will decrease the  186 Os value of the system 906 by 0.4 ppm, not increase it by 29 ppm, as observed for the Jeesiörova-Kevitsa komatiite 907 source. Hence, similarly to Model 1, Model 2 can account for the Nd, Hf, and 187 Os isotopic 908 compositions of the komatiites, but it also fails for 186 Os. 909 One significant modification to Models 1 and 2, which would make it possible to account 910 for all the observed isotopic and elemental data for the komatiites, is assuming early 911 perturbation of the Pt/Os ratio in the komatiite source, without affecting Sm-Nd and Lu-Hf 912 evolution, as well as not modifying Re-Os evolution outside of the chondritic range. The most 913 parsimonious process that can achieve this perturbation is crystal-liquid fractionation of 914 metal, given that fractional crystallization of metal can strongly fractionate Pt/Os, but only 915 modestly fractionates Re/Os, with solid metal-liquid metal bulk distribution coefficients in 916 this order: D Os >D Re >>D Pt >1 (e.g., [START_REF] Cook | Pt-Re-Os 1100 systematics of Group IIAB and IIIAB iron meteorites[END_REF]. Thus, addition to a mantle domain of 917 evolved metal either through late accretion of a fragment of a differentiated planetesimal core, 918 similar to the mechanism proposed by Fischer-Gödde et al. ( 2011), or addition of metal 919 derived from the Earth's core, but partially crystallized at or near the core-mantle boundary, 920 similar to the mechanism of core-mantle interaction proposed by [START_REF] Humayun | A model for osmium isotopic evolution of metallic solids at the core-1188 mantle boundary[END_REF], can best 921 achieve the Os isotopic requirements for the komatiites. Either process would have the added 922 benefit of decreasing the  182 W value of the mixed domain to a negative value, as suggested 923 by the contamination-corrected  182 W = -10 for the komatiites. 924

Compilation of time-integrated 190 Pt/ 188 Os and 187 Re/ 188 Os ratios for IIAB, IIIAB, and 925 IVA magmatic iron meteorites from [START_REF] Cook | Pt-Re-Os 1100 systematics of Group IIAB and IIIAB iron meteorites[END_REF] and [START_REF] Mccoy | Group 1265 IVA irons: New constraints on the crystallization and cooling history of an asteroidal core 1266 with a complex history[END_REF] reveals large, 926 up to three orders of magnitude, variations in the 190 Pt/ 188 Os ratios that resulted from 927 magmatic solid metal-liquid metal fractionation, accompanied by relatively small, about a 928 factor of 2.5, variations in the 187 Re/ 188 Os ratios in the same magmatic systems. The 186,187 Os 929 isotope systematics of the Jeesiörova-Kevitsa komatiite source can be accounted for by 930 mixing with a number of fractionated iron compositions. We performed mixing calculations 931 for Os, Re, and Pt concentrations, as well as for W isotopic composition and W abundances, 932 results of which are illustrated in Fig. 12. For the meteoritic endmember, we chose a low-Os 933 group IIIAB iron meteorite, Tieraco Creek; similar results are also obtained for other low-Os 934 irons from this group, e.g., Grant and Chupaderos. The mantle source endmember used in the 935 mixing calculations was the BSE. The HSE and W abundance data and W isotopic 936 composition for the endmembers were adopted from [START_REF] Cook | Pt-Re-Os 1100 systematics of Group IIAB and IIIAB iron meteorites[END_REF], [START_REF] Becker | Highly 1063 siderophile element composition of the Earth's primitive upper mantle: Constraints from 1064 new data on peridotite massifs and xenoliths[END_REF], 937 [START_REF] Arevalo | Tungsten geochemistry and implications for 1049 understanding the Earth's interior[END_REF], [START_REF] Wasson | Trapped melt in IIIAB irons; solid/liquid elemental partitioning during 1457 the fractionation of the IIIAB magma[END_REF]Kruijer et al. (2017); these are 938 presented in Table 9. Based on the results of calculations, addition of 0.2% of material with 939 the composition of Tieraco Creek to a mantle domain of BSE composition can explain the 940 HSE and W abundances and Os and W isotopic compositions of the mantle domain that gave 941 rise to the Jeesiörova-Kevitsa komatiite system. One caveat to this model is that the mixing 942 event must have occurred very early, within the first 100 Ma of Solar System history, while 943 late accretion of substantial mass was still possible. 944 Addition of fractionated metal to the Moon (and presumably Earth, too) in the form of a 945 late accretionary component was advocated by Fischer-Gödde and [START_REF] Fischer-Gödde | Osmium isotope and highly siderophile element 1118 constraints on ages and nature of meteoritic components in ancient lunar impact rocks[END_REF] in their 946 study of lunar impact melt breccias to explain suprachondritic Pt/Ir, Ru/Ir, and Pd/Ir ratios in 947 some of the impact melts. It has also been argued that the impactors that created the largest 948 impact basins on the Moon and Mars, and presumably Earth as well, were hundreds of km in 949 diameter [START_REF] Ryder | Mass flux in the ancient Earth-Moon system and benign implications for the 1389 origin of life on Earth[END_REF][START_REF] Strom | The origin of 1414 planetary impactors in the inner Solar system[END_REF][START_REF] Bottke | Stochastic 1078 Late Accretion to Earth, the Moon, and Mars[END_REF] and represented by 950 differentiated planetesimals and, thus, would be capable of generating large HSE and negative 951 182 W heterogeneities in portions of the mantle [START_REF] Marchi | Heterogeneous delivery of silicate and 1257 metal to the Earth by large planetesimals[END_REF][START_REF] Marchi | A compositionally heterogeneous martian 1259 mantle due to late accretion[END_REF]. Cores of 952 differentiated planetesimals have up to two orders of magnitude higher W abundances and up 953 to 350 ppm less radiogenic 182 W/ 184 W than the modern BSE [START_REF] Kleine | Rapid accretion and early core 1219 formation on asteroids and the terrestrial planets from Hf-W chronometry[END_REF][START_REF] Kleine | 182 Hf-182 W isotope 1222 systematics of chondrites, eucrites, and martian meteorites: Chronology of core formation 1223 and early mantle differentiation in Vesta and Mars[END_REF]954 Schoenberg et al., 2002;[START_REF] Yin | A 1477 short timescale for terrestrial planet formation from Hf-W chronometry of meteorites[END_REF]. Thus, any mantle domain to which an excess of this 955 material was added would be 182 W-depleted. Such a mantle domain would also be expected to 956 be enriched in HSE, compared to BSE, particularly in Pt and Pd. 957

In order to further evaluate the effect of addition of such material on 182 W/ 184 W and HSE 958 abundances, we plot the  182 W in the Jeesiörova-Kevitsa komatiite system versus the total 959 calculated HSE abundances in its mantle source relative to those in the estimates for modern 960 BSE (Fig. 13). This proportion corresponds to the fraction of the total HSE budget of the BSE 961 added during late accretion assuming an HSE-free mantle following core formation. The 962 182 W/ 184 W of the BSE prior to late accretion is constrained by mass-balance calculations (e.g., 963 [START_REF] Kleine | Tungsten isotopes in planets[END_REF] and is supported by the 182 W/ 184 W data for the lunar mantle of 964 +25±5 ppm [START_REF] Kruijer | Lunar tungsten isotopic 1233 evidence for the late veneer[END_REF][START_REF] Touboul | Tungsten isotopic evidence for 1433 disproportional late accretion to the Earth and Moon[END_REF]Kruijer and Kleine, 2017). Calculations 965 indicate that the estimated 10 ppm deficit in  182 W is consistent with the calculated total HSE 966 abundances in the source of the Jeesiörova-Kevitsa komatiite system and the results of the 967 mixing calculations presented above. 968

A second type of processes to consider is interaction between Earth's core and mantle. 969 This type of process was originally proposed by [START_REF] Walker | 187 Os enrichment in some plumes: 1444 evidence for core-mantle interaction?[END_REF][START_REF] Walker | Applications of the 190 Pt-186 Os isotope system to geochemistry and 1447 cosmochemistry[END_REF] and Brandon et al. 970 (1998Brandon et al. 970 ( , 1999) ) as a means of generating coupled 186,187 Os enrichments observed in some ocean 971 island basalts. Later, [START_REF] Humayun | A model for osmium isotopic evolution of metallic solids at the core-1188 mantle boundary[END_REF] proposed a mechanism of core-mantle interaction that 972 involved outer core liquid becoming trapped in a cumulate pile of Fe-rich non-metallic 973 precipitates (FeO, FeS, Fe 3 Si) at the top of the core. This metal is presumed to undergo partial 974 fractional crystallization. The resulting, chemically-evolved liquid is then incorporated into 975 the base of the mantle, where mantle plumes subsequently entrain it. If this metal had a 976 composition similar to that of evolved iron meteorites, such as Tieraco Creek, its entrainment 977 into the plume that gave rise to the Jeesiörova-Kevitsa komatiites could explain the 186,187 Os 978 isotope systematics of these komatiites. As with the grainy late accretion model, this 979 mechanism has the added benefit of explaining the W isotope systematics of the komatiites. 980

Based on the discovery of negative  182 W values in modern ocean island basalts (Mundl et 981 al., 2017), [START_REF] Rizo | 182 W evidence for core-mantle 1376 interaction in the source of mantle plumes[END_REF] and [START_REF] Mundl-Petermeier | Anomalous μ 182 W signatures in high 3 He/ 4 He ocean island basaltsfingerprints of 1286 Earth's core?[END_REF] have argued for some form 982 of core-mantle exchange/equilibration in order to transfer the inferred negative  182 W value 983 of the core ( 182 W = -220) to plumes rising from the core-mantle boundary. Although 186 Os-984 182 W data for the same rocks are still limited, all Hawaiian lavas with negative  182 W (ranging 985 between -7.8 and -20.2) are also characterized by positive  186 Os (ranging between +18.6 and 986 +62.8) and  187 Os (ranging between +2.0 and +7.2) values (Fig. 14). The projected negative 987  182 W value for the Jeesiörova-Kevitsa komatiites is within the range of  182 W values for 988

Hawaii, implying that these komatiites could be a Proterozoic equivalent of Hawaiian picrites. 989

This process would be expected to have decreased 182 W/ 184 W in the komatiite source relative 990 to the ambient mantle, and also potentially have increased the HSE abundances over the 991 ambient mantle levels, which is consistent with the observations. 992

We conclude that at present, both grainy late accretion and core-mantle interaction remain 993 viable mechanisms for explaining the observed combined lithophile trace-and siderophile 994 element isotope and abundance systematics of the Jeesiörova-Kevitsa komatiite system. 995 996

Conclusions 997

The Jeesiörova-Kevitsa komatiites of the Fennoscandian Shield are unique among the 998 komatiite systems studied so far. They have a relatively young age of 2.05 Ga that has not 999 been reported for the existing terrestrial komatiite record elsewhere, yet are as MgO-rich as 1000 some of their late Archean counterparts. These komatiites were derived from a contemporary 1001 DMM-like mantle source in terms of Nd-Hf isotope systematics; this source was 1002 characterized by a negative 182 W anomaly, total absolute HSE abundances ~20% higher than 1003 of the modern BSE, and chondritic initial 187 Os/ 188 Os, but suprachondritic initial 186 Os/ 188 Os 1004 isotopic compositions. When considered together, these constraints are reconciled within a 1005 model that requires derivation of the parental Jeesiörova-Kevitsa komatiite magma from a 1006 mantle domain that experienced a melting and melt depletion event within the first 100 Ma of 1007 Solar System history. This mantle domain is argued here to have acquired its HSE and Os 1008 isotope systematics as a result of either (1) grainy late accretion of differentiated metallic 1009 impactors or (2) core-mantle interaction via addition of chemically fractionated outer core 1010 metal to the mantle source domain of the komatiites. The survival of the early-formed 1011 Jeesiörova-Kevitsa komatiite mantle source for ≥2.5 billion years implies that portions of the 1012 mantle remained poorly mixed with respect to HSE and W until at least the Paleoproterozoic. 1013 1014 melting. The normalizing value for W is from [START_REF] Arevalo | Tungsten geochemistry and implications for 1049 understanding the Earth's interior[END_REF] and, for 1506 the other elements, from [START_REF] Hofmann | Chemical differentiation of the Earth: The relationship between 1170 mantle, continental crust and oceanic crust[END_REF]. 1507 and Paleoproterozoic komatiite systems plotted as a function of age. The data are from 1527 [START_REF] Foster | Re-Os isotopic evidence for 1130 genesis of Archaean nickel ores from uncontaminated komatiites[END_REF]; Puchtel et al. (2004a;[START_REF] Mutanen | The Akanvaara intrusion and the Keivitsa-Satovaara Complex, with stops 1293 at Kaikkivaltiaanlehto and Särkivaara intrusions[END_REF]2007;2009a;2009b;2014;2016a;2016b;15282018), and this study. (2004a, 2005, 2009b, 2014, 2016b; 2018), and this study. The data for chondritic 1533 meteorites are compiled from [START_REF] Horan | Highly 1178 siderophile elements in chondrites[END_REF], [START_REF] Brandon | Re-Os isotopic systematics 1087 and platinum group element composition of the Tagish Lake carbonaceous chondrite[END_REF][START_REF] Brandon | Platinum-osmium isotope evolution of 1090 the Earth's mantle: Constraints from chondrites and Os-rich alloys[END_REF] Data (in wt.% except for Cr in ppm) were recalculated on an anhydrous basis, but not re-normalized to 100% in order to preserve information on the quality of the analyses. Here and in Tables 2 and3, samples from 12D-PPR to LP10 are Jeesiörova komatiites and those from KD03 to KD14 -Kevitsa komatiite dikes. Abundances recalculated on an anhydrous basis in ppm, except for W (ppb). N-normalized to the BSE values of [START_REF] Hofmann | Chemical differentiation of the Earth: The relationship between 1170 mantle, continental crust and oceanic crust[END_REF] for all elements except Ga [START_REF] Mcdonough | The composition of the Earth[END_REF] and W [START_REF] Arevalo | Tungsten geochemistry and implications for 1049 understanding the Earth's interior[END_REF], as reported in the table. HSE abundances recalculated on an anhydrous basis. Initial  187 Os values were calculated at T = 2049 Ma derived from the isochron using the parameters specified in the text. The HSE normalizing values are from [START_REF] Horan | Highly 1178 siderophile elements in chondrites[END_REF]. Replicatereplicate digestion of the sample. Chromite separates 12D-PPR_1 through 12D-PPR_6 and 13-EJH_1 and 13-EJH_2 denote spiked subsamples of unspiked digestions for the Pt-Os study.

Figure captions 1484

  202 4.1. Analysis of mineral compositions 203 of the measured 143 Nd/ 144 Nd ratio for both samples and standards was between 5 and 7 ppm (2SE). During the 281 course of the present analytical campaign, the external reproducibility of the 60 ppb AMES Nd standard solution 282 measurements for 143 Nd/ 144 Nd was ±7.0 ppm (2SD). The average 143 Nd/ 144 Nd measured for the AMES Nd 283 standard during each analytical session was used to calculate the instrumental mass bias coefficient for 284 correction of the measured 143 Nd/ 144 Nd in the samples using the long-term average 143 Nd/ 144 Nd value measured 285

  compositions and Lu and Hf concentrations in the Jeesiörova komatiites were determined at 299 the IGL and the Department of Terrestrial Magnatism (DTM), Carnegie Institution for Science. At the IGL, 300 approximately 500 mg of the whole-rock sample powder and pure clinopyroxene separate were weighted out in 301 15 mL Teflon inserts of Parr bombs with 7 mL double-distilled concentrated HF, 1.5 mL HNO 3 , and an 302 appropriate amount of mixed 176 Lu-178 Hf spike, sealed in stainless steel jackets and digested in an oven at 170°C

  on the 187 Re/ 188 Os versus 443 187 Os/ 188 Os diagram in Fig. 6, on a CI chondrite-normalized plot in Fig. 7, and on MgO 444 versus HSE variation diagrams in Fig. 8. 445

  on the 478 190 Pt/ 188 Os versus 186 Os/ 188 Os diagram in Fig. 6B. The data for the seven chromite separates 479 define a regression line with a slope corresponding to an age of ~2.0 Ga and an average initial 480 186 Os/ 188 Os = 0.1198369±3 ( 186 Os = +29±2, 2SE) calculated for the time of the komatiite 481 emplacement derived from the Re-Os isochron (2049 Ma). 482 483

  and plotted in Fig. 4B and as BSE-normalized values in Fig. 5. The 486 W abundances vary between 8.8 and 13 ppb in the Kevitsa komatiites; the only analyzed 487 sample of the Jeesiörova komatiite contains 27 ppb W. In the MgO versus W variation 488 diagram (Fig. 4B), the data for the Kevitsa komatiites plot with some scatter around a nearly 489 horizontal trend line, whereas the only data point for the Jeesiörova komatiite plots above the 490 trend line. 491 The BSE-normalized trace element abundances (Fig. 5) for the Kevitsa komatiites are 492 characterized by small negative to positive W abundance offsets relative to neighboring 493 elements (i.e., Th and U) with similar incompatibility during mantle melting (W/W* = 494 0.95±0.13, where W/W* = W N /(√[Th N ×U N ]) and N are BSE-normalized values from Arevalo 495 and McDonough (2008) and Hofmann (1988)). The W/W* ratio in the only Jeesiörova 496 komatiite sample analyzed (LP10) is 2.1. 497

  and had long-term coupled Sm/Nd and Lu/Hf isotope 780 systematics (Blichert-Toft and Puchtel, 2010). If the source had originally chondritic Sm/Nd 781 and Lu/Hf, then it experienced a melt-extraction event a minimum of several hundred million 782 years prior to komatiite emplacement. The same melt extraction event might also have been 783 expected to result in depletion of incompatible Re relative to compatible Os, as well as more 784 modest depletion in Pt relative to Os. Over time, these chemical effects would have led to the 785 development of negative  187 Os and  186 Os values. Instead, as shown in Section 6.3, the 786

Fig. 1 .Fig. 2 .Fig. 3 .Fig. 4 .Fig. 5 .

 12345 Fig. 1. (a) Geological map showing the distribution of the 2.05 Ga komatiites and related 1485 picrites in the Central Lapland Greenstone Belt, Finland (modified after Hanski et al., 1486 2001), and the locations of the Jeesiörova komatiites and the Kevitsa mafic-ultramafic 1487 intrusion. (b) Geological map of the Kevitsa intrusion (modified after Mutanen and 1488 Huhma, 2001). (c) Location of drill hole DDH 814 and other drill holes penetrating 1489 Kevitsa komatiite dikes (red dots). 1490

Fig. 6 .Fig. 7 .Fig. 8 .

 678 Fig. 6. (a) Re-Os isochron diagram for whole-rock samples and olivine and chromite 1508 separates from the Jeesiörova and Kevitsa komatiites analyzed in this study. (b) Pt-Os 1509 isotopic data for chromite separates from the Jeesiörova komatiites. 1510 1511 Fig. 7. CI chondrite-normalized HSE abundances in whole-rock samples of the (a) Jeesiörova 1512 komatiites, (b) Kevitsa komatiites, and (c) olivine separates from the Kevitsa komatiites. 1513

Fig. 9 .

 9 Fig. 9. (a) 147 Sm-143 Nd and (b) 176 Lu-176 Hf isochron diagrams for the Jeesiörova-Kevitsa 1516 komatiites and clinopyroxene separates. See text for additional details. (c) Diagram 1517 illustrating average initial  143 Nd and  176 Hf values for the emplaced Jeesiörova-Kevitsa 1518 komatiites (light-green symbol) and their mantle source (dark-green symbol), as well as 1519 for the Chondritic Uniform Reservoir (CHUR) and Depleted MORB Mantle (DMM) 1520 calculated at the time of the emplacement (2049 Ma). The CHUR and DMM parameters 1521

Fig

  Fig. 1. 1614 1615

. 385 5.2.1. Olivine. Olivine

	386 387 391 392	exhibits a continuous compositional range from Fo 94.1 to Fo 73.8 , with the highest forsterite content reported for large phenocrysts and the lowest for small are considered to be cognate phases crystallized from the same komatiitic parental magma. grains in the groundmass. Microphenocrysts in the chilled margin samples vary from Fo 83.5 to absence of xenocrystic mantle olivine among the grains analyzed. Instead, all olivine grains 5

.2.2. Pyroxenes and plagioclase. Groundmass clinopyroxene

	393	has an augitic average
	394	composition of En 49 Fs 10 Wo 41 with Al 2 O 3 ranging from 2.4 to 5.8 wt.% and TiO 2 from 0.12 to
	395	0.71 wt.%. Orthopyroxene grains average En 71 Fs 25 Wo 4 and have 3.4-4.5 wt.% Al 2 O 3 and
	396	0.40-0.57 wt.% TiO 2 . The anorthite content of the analyzed plagioclase grains in the
	397	groundmass ranges from 57.7 to 72.1 mol.%.
	398	5.2.3. Chrome spinel. The most primitive chrome spinel grains have MgO contents of
	399	around 15-16 wt.% and corresponding Mg# of 0.71-0.76, whereas in the most evolved
	400	grains, MgO and Mg# drop to 2.7 wt.% and 0.14, respectively. The MgO-rich grains show an
	401	average Al 2 O 3 content of 14.7 wt.% and an average TiO 2 concentration of 0.38 wt.%, yielding
	402	Al 2 O 3 /TiO 2 typical of well-preserved chrome spinel grains in komatiitic lavas in Central

  The data for chondritic meteorites are compiled fromWalker et al. 1529[START_REF] Walker | 1452 Comparative 187 Re-187 Os systematics of chondrites: Implications regarding early solar 1453 system processes[END_REF],[START_REF] Brandon | Re-Os isotopic systematics 1087 and platinum group element composition of the Tagish Lake carbonaceous chondrite[END_REF], and Fischer-Gödde et al. (2010). (b). Initial 186 Os/ 188 Os 1530 isotopic compositions, expressed in terms of  186 Os, of Archean and Paleoproterozoic 1531 komatiite systems studied to date, plotted as a function of age. The data are from Puchtel 1532 et al.

Table 1 .

 1 Calculated total Pt and Pd abundances in the sources of Archean and 1536 Paleoproterozoic komatiite systems plotted as per cent of the total Pt and Pd abundances 1537 in the estimates for the modern BSE of Becker et al. (2006). Hf isotope systematics of the Jeesiörova-Kevitsa komatiites. The data for the 1550 IIIAB iron meteorite Tieraco Creek used in the modeling are from Cook et al. (2004). 1551 Addition of 0.2% of Tieraco Creek metal to the mantle domain with BSE Re/Os and Pt/Os 1552 prior to the global melt depletion event can explain the 186,187 Os/ 188 Os isotope systematics 1553 of the Jeesiörova-Kevitsa komatiite source. See text for more details. 1554 Major and minor element data for Jeesiörova and Kevitsa komatiites.

	, and

182

W value of the BSE prior to late 1559 accretion is constrained via mass-balance calculations and is supported by the W isotopic 1560 data for the lunar mantle ( 182 W = +25±5 ppm:

[START_REF] Kruijer | Lunar tungsten isotopic 1233 evidence for the late veneer[END_REF][START_REF] Touboul | Tungsten isotopic evidence for 1433 disproportional late accretion to the Earth and Moon[END_REF]

. 1561

Table 3 .

 3 Highly siderophile element abundances (in ppb) and Re-Os isotopic data for Jeesiörova and Kevitsa komatiites and olivine and chromite separates.

	Sample No.	Re	Os	Ir	Ru	Pt	Pd	187 Re/ 188 Os	187 Os/ 188 Os	 187 Os(2049)
	12D-PPR WR	0.9867	5.434	4.226	7.436	17.89	15.35	0.8766±0.0035	0.14236±7	-1.0
	13-EJH WR	0.6650	4.269	3.279	6.235	18.26	14.87	0.7517±0.0030	0.13836±6	-0.7
	17.1-PPR WR	0.8497	3.213	2.637	5.869	19.17	15.22	1.280±0.005	0.15817±7	+0.6
	LP-10 WR	0.3756	4.435	3.631	6.657	16.75	13.93	0.4082±0.0016	0.12871±6	+1.3
	KD-03 WR	1.188	1.528	1.472	2.573	22.31	19.01	3.805±0.015	0.24505±2	-0.1
	KD-05 WR	1.007	2.149	1.845	3.549	19.77	16.92	2.277±0.009	0.19117±8	-0.9
	KD-06 WR	1.010	5.528	3.006	3.824	19.57	16.49	0.8817±0.0035	0.14353±7	-0.1
	KD-07 WR	0.8141	4.161	2.571	4.082	16.82	14.72	0.9452±0.0038	0.14662±1	+0.7
	KD-08 WR	0.8693	3.684	2.824	4.300	17.77	15.32	1.141±0.005	0.15200±8	-0.6
	KD-09 WR	0.7062	3.956	2.870	5.001	16.02	13.74	0.8620±0.0035	0.14307±6	+0.1
	KD-10 WR	0.6813	3.190	2.434	5.027	16.22	13.60	1.032±0.004	0.14820±7	-0.6
	KD-11 WR	0.9146	3.701	2.708	4.262	19.20	15.87	1.195±0.005	0.15471±7	+0.1
	KD-12 WR	1.051	2.297	1.753	3.545	20.43	17.04	2.223±0.009	0.19087±10	+0.5
	KD-13 WR	1.127	3.456	2.166	3.436	21.10	17.97	1.579±0.006	0.16626±10	-1.4
	Replicate	1.109	2.063	1.933	3.384	21.04	17.76	2.617±0.010	0.20361±9	-0.3
	KD-14 WR	1.284	2.035	1.593	2.168	23.68	20.01	3.076±0.012	0.21969±2	-0.2
	12D-PPR Chr	1.195	122.7					0.0469±0.0004	0.11378±5	-0.8
	17.1-PPR Chr	1.233	11.13					0.5341±0.0044	0.13226±8	+0.6
	12D-PPR_1 Chr	1.688	156.7	102.2		27.85		0.0518±0.0010	0.11462±7	-0.2
	12D-PPR_2 Chr	0.8898	141.0	94.4		33.03		0.0304±0.0010	0.11357±6	-0.5
	12D-PPR_3 Chr	0.6325	149.8	102.5		36.05		0.0203±0.0009	0.11336±6	-0.4
	12D-PPR_4 Chr	0.5442	153.3	103.6		38.92		0.0171±0.0009	0.11339±6	-0.2
	12D-PPR_5 Chr	0.6303	150.2	101.7		47.32		0.0202±0.0009	0.11329±7	-0.4
	12D-PPR_6 Chr	0.5624	157.5	108.0		52.97		0.0172±0.0009	0.11337±5	-0.3
	13-EJH_1 Chr	0.2597	53.32	38.33		6.164		0.0234±0.0027	0.11394±6	+0.1
	13-EJH_2 Chr	0.3449	53.59	38.45		7.314		0.0310±0.0026	0.11393±7	-0.2
	KD-05 Ol	0.04468	3.952	2.262	2.296	1.525	1.426	0.0544±0.0009	0.11455±4	-0.4
	KD-10 Ol	0.6178	3.126	2.285	4.158	1.936	1.698	0.9547±0.0050	0.14627±4	0.0
	KD-05 Chr	0.4528	35.29					0.0617±0.0016	0.11497±6	-0.2
	KD-08 Chr	0.5359	15.54					0.1660±0.0023	0.11837±8	-0.4
	KD-10 Chr	1.197	78.14					0.0737±0.0011	0.11531±7	-0.3
	KD-12 Chr	0.7288	7.693					0.4568±0.0062	0.12945±13	0.0
	LP-10 Chr	1.014	55.28					0.0883±0.0019	0.11655±7	-0.2

Table 4 .

 4 High-precision Os isotopic data and Re/Os and Pt/Os ratios for chromite separates from the Jeesiörova komatiites.187 Os and  186 Os values were calculated at T = 2049 Ma using the parameters specified in the text.

	Sample No.	187 Re/ 188 Os	187 Os/ 188 Os	 187 Os(T)	190 Pt/ 188 Os	184 Os/ 188 Os	186 Os/ 188 Os	 186 Os(T)
	12D-PPR_1 Chr 0.0518±0.0010 0.1146036±06	-0.23	0.0001697±34 0.0013031±2 0.1198369±6	+25±5
	12D-PPR_2 Chr 0.0304±0.0010 0.1135064±05	-0.54	0.0002238±45 0.0013028±2 0.1198381±5	+34±4
	12D-PPR_3 Chr 0.0203±0.0009 0.1134565±07	-0.27	0.0002299±46 0.0013023±3 0.1198377±7	+30±6
	12D-PPR_4 Chr 0.0171±0.0009 0.1134060±05	-0.22	0.0002425±48 0.0013033±2 0.1198381±5	+33±4
	12D-PPR_5 Chr 0.0202±0.0009 0.1134294±06	-0.29	0.0003009±60 0.0013033±2 0.1198378±6	+29±5
	12D-PPR_6 Chr 0.0172±0.0009 0.1134048±06	-0.22	0.0003213±64 0.0013033±2 0.1198376±6	+27±5
	13-EJH Chr	0.0272±0.0019 0.1138779±11	-0.11	0.0001204±24 0.0013024±4 0.1198370±10	+27±8
	Initial 

Table 6 .

 6 Sm-Nd isotope and concentration data for Jeesiörova and Kevitsa komatiites. 1588 Initial  143 Nd values were calculated at T = 2049 Ma using the parameters specified in the 1589 text. WR -whole-rock samples; Cpx -clinopyroxene separates.

	Sample	Sm, ppm Nd, ppm	147 Sm/ 144 Nd	±2SD	143 Nd/ 144 Nd	±2SE	 143 Nd(T)
	13-EJH WR	1.081	2.446	0.2673	0.0005	0.513768	0.000004	+3.5
	17.1-PPR WR	1.123	2.445	0.2779	0.0006	0.513917	0.000003	+3.6
	KD-05 WR	1.379	2.966	0.2811	0.0006	0.513959	0.000004	+3.6
	KD-08 WR	1.176	2.546	0.2794	0.0006	0.513950	0.000003	+3.8
	KD-10 WR	0.9099	1.950	0.2821	0.0006	0.513981	0.000003	+3.7
	KD-12 WR	1.501	3.245	0.2798	0.0006	0.513946	0.000003	+3.7
	13-EJH Cpx	0.9271	1.542	0.3636	0.0007	0.515069	0.000003	+3.5
	17.1-PPR Cpx	0.9256	1.496	0.3743	0.0008	0.515217	0.000003	+3.6
	KD-05 Cpx	0.6775	1.309	0.3131	0.0006	0.514395	0.000003	+3.7
	KD-08 Cpx	0.6564	1.335	0.2975	0.0006	0.514177	0.000004	+3.5
	KD-10 Cpx_1	0.6936	1.368	0.3066	0.0006	0.514317	0.000004	+3.9
	KD-10 Cpx_2	0.7036	1.388	0.3066	0.0006	0.514316	0.000003	+3.8
	KD-12 Cpx	0.8599	1.676	0.3104	0.0006	0.514366	0.000003	+3.8
	1590
	1591

Fo 93.2 . The CaO content of all analyzed olivine grains is 0.20-0.32 wt.%, typical of magmatic

olivine. This, together with a negative correlation between MgO and MnO, indicates the

Acknowledgements 1015

This study was supported by NSF Petrology and Geochemistry Grants EAR 1447174 and 1016 EAR 1754186 to ISP, EAR-1624587 to RJW and AMP, and Academy of Finland Grant 1017 281859 to EJH. AMP would like to acknowledge support from FWF grant V659-N29. We 1018 thank V.A. Puchtel and R.W. Nicklas for help with preparation of samples for trace element 1019 analysis and ICP-MS measurements. Thorough and constructive reviews by Klaus Mezger, 1020 David van Acken, and an anonymous reviewer helped improve the original version of the 1021 manuscript. We thank Catherine Chauvel for comments and editorial handling. 1022 1023

The composition of the Jeesiörova-Kevitsa komatiite mantle source is calculated to have 1607 resulted from mixing of 0.2% of material of the IIIAB iron meteorite Tieraco Creek into a 1608 mantle domain with the BSE composition. Uncertainties on the HSE abundances in the 1609 average estimate for the BSE are 1SE [START_REF] Becker | Highly 1063 siderophile element composition of the Earth's primitive upper mantle: Constraints from 1064 new data on peridotite massifs and xenoliths[END_REF]. 1610 A1 andA2 Click here to download Background dataset for online publication only: Supplementary Tables A1 andA2.docx Supplementary Table A3 Click here to download Background dataset for online publication only: Supplementary Table A3.xlsx

Declaration of interests

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.