The first major rise in atmospheric oxygen levels ~2.4-2.2 Ga, the Great Oxidation Event (GOE), has been invoked as a driver for increased concentrations of redox-sensitive elements in rocks that formed in Earth's surface environments (e.g., (1)(2)(3)(4)(5)(6)). The GOE was a protracted shift in atmospheric oxygen levels, though the architecture of atmospheric O2 oscillations in its wake is uncertain (e.g., [7][8][9][10]. Preservation of detrital uraninite and pyrite in fluvial sediments, as well as sulfur isotope mass-independent fractionation recorded in sulfides in shallow marine systems, is commonly interpreted as evidence for atmospheric O2 concentrations <10 -5 the present atmospheric level before the GOE (e.g., (2,(11)(12)(13)(14)). Additionally, a shift from reduced to oxidized paleosols is documented between 2.4 and 2.2 Ga (15,16), and marine shales record pronounced enrichments of redox-sensitive metals after the GOE (e.g., (17,18)). During the GOE, ~2.3-2.0 Ga marine carbonates record positive δ 13 C values, some higher than +10‰, in sharp contrast with the near-0‰ δ 13 C values that characterize marine carbonates immediately before the GOE and for most of Earth history (19). This global positive δ 13 C excursion, the Lomagundi-Jatuli Event (LJE), has been interpreted to record a period of extensive organic carbon burial, which would have produced a substantial flux of O2 to the atmosphere (8,19,20), though the feasibility of such a global-scale event has been questioned (21)(22)(23)(24)(25). The hypothesis of substantial carbon burial, however, is consistent with evidence for high surface oxidant concentrations during this period (17,18,26,27). At the end of the LJE, it has been suggested that atmospheric O2 levels crashed (8,10,17), though the timing and magnitude of this shift is not well understood. Consequently, many of the details associated with the LJE remain elusive, including the processes related to its initiation and termination (e.g., (19,20)), as well as the duration and intensity of atmospheric oxidation (e.g., (18)).
Although there are oxidized paleosols and rocks that were exposed and weathered during the LJE (e.g., (15,28)), the timing of their oxidation has not been well established. Direct age determination of oxidation products such as Fe(III)-hydroxides/oxides has not been attempted for LJE rocks, but instead the age of oxidation has been indirectly estimated through studies of associated minerals and weathering products (29)(30)(31)(32). In addition, all geologic units that represent the LJE have seen complex tectonothermal histories, and it is possible that redoxsensitive elements have been mobile over this history. Across the Fennoscandian Shield in Arctic Russia, ~2.06 Ga subaerial volcanic rocks of the Kuetsjärvi Volcanic Formation (KVF) have witnessed a complex sequence of post-depositional tectonic and metamorphic events, and it is unclear when the rocks attained their unusually high Fe contents and high ratios of Fe 3+ to total Fe.
Application of the U-Th-Pb isotope system to oxide-rich rocks of Archean age has been used to constrain Archean seawater U contents and evaluate the timing of post-depositional oxidation (33,34). Under the relatively high oxygen conditions that likely characterized the LJE, oxidation of insoluble tetravalent U to soluble hexavalent U would mobilize U, whereas insoluble tetravalent Th would remain immobile (35,36). This in turn would decouple accrual of uranogenic ( 206 Pb, 207 Pb) and thorogenic ( 208 Pb) radiogenic daughter Pb isotopes. Decoupling of these isotope systems can be especially well established in rocks such as mantle-derived basalts, where the primary Th/U ratios are tightly constrained (37)(38)(39)(40)(41), and models of the Pb isotope evolution of the mantle and crust provide a solid framework for interpreting Pb isotope compositions (42,43). Despite their comparatively low U contents, basalts are attractive for tracing U mobility because U is contained in weatherable components such as glass and along grain boundaries (e.g., (44,45)), whereas a significant portion of the U budget of more evolved rocks such as granites is locked up in accessory phases such as zircon, which is resistant to weathering.
Our focus is on oxidative weathering and alteration of basaltic rocks from the KVF, where their high Fe contents provide a sensitive tracer of oxidation of igneous Fe 2+ to Fe 3+ . Application of the U-Th-Pb isotope system to directly date surface redox conditions via the oxidative mobility of U in basalts (33) has been underexplored. Here we demonstrate that the differential mobility of U and Th that occurs upon oxidation of U 4+ to U 6+ in the KVF may be traced using the U-Th-Pb isotope system, confirming that the age of oxidation is shortly after deposition ~2.06 Ga. We suggest that loss of mobile U 6+ was likely a widespread phenomenon on the continents, producing a flux of U to the oceans that eventually found its way to the mantle via subduction, correlating with major shifts in mantle geochemistry. The U-Th-Pb geochronological results show how post-eruptive oxidative U loss may be distinguished from later reintroduction of U during metamorphism, complexities that cannot be understood through simple elemental analysis.
The Fennoscandian Arctic Russia Drilling Early Earth Project (FAR-DEEP) targeted the Pechenga Greenstone Belt, the Imandra-Varzuga Greenstone Belt, and the Onega Basin in a composite collection of scientific drill cores spanning the GOE (46). We focus here on the Pechenga Greenstone Belt (PGB), a thick succession of supracrustal rocks that overlie Archean basement (SI Appendix, Fig. S1; (47)). The uppermost part of this stratigraphy is the Pilgujärvi Volcanic Formation, deposited ~1.9 Ga (48). The Northern Zone of the PGB is composed of four sedimentary-volcanic stratigraphic successions which were sampled by FAR-DEEP (49). The KVF marks the top of the second sedimentary-volcanic sequence and is the most intensely oxidized volcanic formation observed within the Fennoscandian Shield (SI Appendix, Fig. S2; (49,50)). The base of the KVF has a transitional contact with the underlying sedimentary formation and is divided into two nearly equal parts by a 30m-thick conglomerate member. The top of the KVF is marked by an unconformity that contains a modestly developed, oxidized weathering profile (50). Rounded, oxidized KVF clasts are observed within the conglomerate beds of the overlying Kolosjoki Sedimentary Formation, suggesting that oxidation of the KVF occurred prior to deposition of the overlying sedimentary-volcanic succession (28,(49)(50)(51). U-Pb zircon geochronology from the overlying Kolosjoki Sedimentary Formation constrains the maximum age of the Kolosjoki Sedimentary Formation as 2058 ± 6 Ma (52). The geochronology of this sequence is permissive of an age between 2077 and 2052 Ma for the KVF, however, this range is insignificant for our purpose and we take the KVF eruptive age as ~2.06 Ga. Following eruption, the KVF was metamorphosed to lower greenschist facies (52), and this most likely occurred during the 1.92 -1.79 Ga Svecofennian orogeny (53,54).
The samples analyzed in this study (SI Appendix) are from FAR-DEEP core 6A, which intersects the middle subunit of the KVF (51). The KVF in core 6A contains a wide range of compositions, from picrites to felsic lavas and breccias, and is interpreted to have been erupted in an intraplate continental rift (51,55). Many of the lava flows contain abundant amygdales which is taken as evidence for eruption in a subaerial environment, and these are filled with a variety of secondary minerals including adularia, allanite, biotite, calcite, chlorite, epidote, hematite, jasper, phlogopite, quartz, sericite, titanite, elemental sulfur, and axinite (SI Appendix, Fig. S3; (50)). In particular, hematite in the amygdales is interpreted to reflect oxidation and precipitation of deeper Fe-bearing groundwaters upon intersection with more shallow oxygenated fluids during early burial (e.g., (50)). In addition, allanite deposition in the amygdales requires transport of REEs, U, and Th in solution, important to consider in our study, and comparison of Sm-Nd isotope evolution trends of core 6A samples with an allanite-rich sample produce an average allanite formation age that overlaps the depositional age (SI Appendix, Table S1).
Approximately half of the core 6A section is composed of mafic volcanic units or sills. This is corroborated by Zr/Ti ratios (SI Appendix, Fig. S4; Table S2), which may be used as a proxy for the degree of differentiation. All of the mafic units in core 6A have Fe contents that are higher than those expected for mantle-derived melts, and nearly all have higher fractions of Fe 3+ than expected for melts in equilibrium with the mantle (SI Appendix, Fig. S5). Our sampling target was the lower part of Unit E in core 6A, a mafic unit that has exceptionally high Fe contents and high extent of oxidation. High levels of Fe enrichment and oxidation are associated with relative SiO2 enrichment (SI Appendix, Fig. S6), and depletion in MgO and CaO, the latter of which corresponds to relative enrichment of Al2O3 (SI Appendix, Fig. S7). These chemical relations are explained by breakdown of Fe-Mg silicate minerals and feldspar (SI Appendix) during weathering and alteration. Indices of mafic alteration, such as MIAo-K (56), correlate with X-Fe 3+ (50) (Fig. S8; Table S2), implying breakdown of mafic phases as a potential source of oxidized Fe during weathering or fluid-rock interaction. Some of the enrichment of Fe in the mafic units, relative to primary Fe contents expected for mantle-derived magmas, may be explained by crystal fractionation of parental magmas and/or weathering/alteration reactions where Fe released through breakdown of Fe-Mg silicate minerals was retained via in situ oxidation (SI Appendix, Fig. S9). For Unit E samples, however, these processes cannot explain the very high measured Fe contents, which suggests possible addition of Fe from hydrothermal fluids that circulated during early burial.
Nine mafic volcanic samples from the lower part of Unit E were taken for U-Th-Pb isotope analysis, guided by geochemical and petrographic observations that indicated high extents of Fe enrichment and oxidation but relatively low Th contents, suggesting minimal complications from the presence of allanite in amygdales (50). In addition, to test for the possibility that the samples were open to Pb addition during hydrothermal fluid-rock interaction, we analyzed ten Ni-Cu ore samples for their Pb isotope compositions from the 1980 Ma Pilgujärvi Sedimentary Formation, which contains several Ni-Cu ore horizons, commonly within the contact zones of organic-rich shales and mafic/ultramafic (ferropicrites) igneous bodies (57).
Uranium, Th, and Pb concentrations of the Kuetsjärvi lavas (0.57 to 0.96 ppm U, 1.30 to 4.21 ppm Th, 3.41 to 7.13 ppm Pb) are consistently low and overlap those expected for basaltic, mantle-derived magmas (58), with the exception of the sample at 97.52 m, which contains 2.82 ppm U, 24.28 ppm Th, and 8.20 ppm Pb (Table 1). The majority of mafic rocks from core 6A, and Unit E, have low Th/Zr ratios, indicating derivation from a common parental magma and minimal influence from weathering on relative Th and Zr concentrations (SI Appendix). In contrast, the sample at 97.52 m has an anomalously high Th/Zr ratio, confirming the influence of post-eruptive allanite addition. U/Zr ratios for mafic samples in core 6A are more variable than Th/Zr (SI Appendix, Fig. S10), suggesting post-depositional U mobility, and this is also indicated by the relatively large range in Th/U ratios relative to the narrow range expected for primary igneous compositions (SI Appendix). Broadly, samples from Unit E scatter about the crustal average Th/U ratio, with the exception of the samples at 97. 52 and 97.36 m. Whole-rock 206 Pb/ 204 Pb, 207 Pb/ 204 Pb, and 208 Pb/ 204 Pb ratios of the Kuetsjärvi lavas are 17.20-21.28, 15.25-15.67, and 38.91-52.44, respectively (Table 1). These compositions distribute along a linear 206 Pb/ 204 Pb-207 Pb/ 204 Pb array (SI Appendix, Fig. S11) with a slope equivalent to a date of 1711 ± 130 Ma (2SE, MSWD = 3.3). In contrast, the 238 U/ 204 Pb-206 Pb/ 204 Pb, 235 U/ 204 Pb-207 Pb/ 204 Pb, and 232 Th/ 204 Pb-208 Pb/ 204 Pb isochron dates are all significantly younger (SI Appendix, S3). Ore samples from the ca.
1980 Ma Pilgujärvi Volcanic Formation have low 206 Pb/ 204 Pb, 207 Pb/ 204 Pb, and 208 Pb/ 204 Pb that vary within a relatively tight range of 15.21-17.18, 15.14-15.44, and 34.82-36.43 (SI Appendix, Table S4), respectively, consistent with the Ni-Cu ore measurements of Hanski (1992).
U-Th-Pb geochronology can provide time information on the differential redox behavior of U relative to Th, making it an ideal system for assessing the redox history of the Kuetsjärvi lavas. The 206 Pb/ 204 Pb-207 Pb/ 204 Pb regression using all data, including leaches, produces a date of 1770 ± 71 Ma (2SE, MSWD = 7.7; Fig. 1). The absence of correlation between 208 Pb/ 204 Pb, 207 Pb/ 204 Pb, and 206 Pb/ 204 Pb with 1/Pb in these rocks, coupled with the absence of mixing lines with the Pb isotope compositions of local Pilgujärvi Ni-Cu ores (SI Appendix, Fig. S14, Table S4), indicates that the 206 Pb/ 204 Pb-207 Pb/ 204 Pb array does not represent mixing with an overprinting Pb reservoir. While Pb addition is thus unlikely, the difference between the 207 Pb/ 204 Pb-206 Pb/ 204 Pb age of ~1.8 Ga and the variable 238 U/ 204 Pb-206 Pb/ 204 Pb, 235 U/ 204 Pb-207 Pb/ 204 Pb, and 232 Th/ 204 Pb-208 Pb/ 204 Pb isochron dates (~1.2, 1.4, and 1.3 Ga, respectively) is most aptly modeled by a Pb-loss event between ~200 and ~400 Ma (SI Appendix, Fig. S15), the timing of which aligns with the ~250-350 Ma tectonothermal event in the Kola Peninsula region documented by apatite fission-track thermochronology (59). This Pb-loss event did not result in any significant change in isotopic composition through mass-dependent fractionation and had a minimal effect on the 206 Pb/ 204 Pb-207 Pb/ 204 Pb age (Fig. 1; SI Appendix). We therefore ascribe age significance to the ~1.8 Ga linear 207 Pb/ 204 Pb-206 Pb/ 204 Pb array of these lavas. This age is significantly younger than the ~2.06 Ga eruption age ( 52), and appears coincident with the Svecofennian Orogeny (53,54).
Three key features of the Pb isotope compositions of the Kuetsjärvi lavas document early mobilization of U shortly after eruption at 2.06 Ga, followed by reintroduction of U at ~1.8 Ga: 1) a 207 Pb-206 Pb age of ~1.8 Ga (Fig. 1b); 2) when compared to the evolution of average crust and mantle as represented by the Stacey-Kramers evolution curve (S-K; (36)), the rocks bear elevated thorogenic Pb ( 208 Pb) relative to uranogenic Pb ( 206 Pb, 207 Pb) (Figs. 1a, 2a); and 3) the uranogenic Pb is substantially unradiogenic relative to 2.06-1.80 Ga mantle and crust (Figs. 1b, 2b), producing an intersection with the S-K curve at an age older than the eruption age, and markedly older than the 207 Pb-206 Pb age, causing the array to plot significantly below the modern endpoint of the S-K curve. Such relations can only be explained by a major, early disruption in U-Th-Pb isotope evolution, analogous to the inferences of U-Th decoupling proposed by whole-rock U-Pb isotope studies of high-grade metamorphic rocks where U was removed relative to Th (SI Appendix, Table S5; (60)).
The offset between thorogenic and uranogenic Pb in the KVF is evidence for substantial, perhaps complete, removal of uranium from these rocks shortly after deposition. In Figure 2, the stepwise evolution of the U-Th-Pb isotope compositions is depicted from 2.06 Ga to the present, charting how the Pb isotope compositions reflect multiple post-eruptive processes. At 1.8 Ga, the time of 207 Pb/ 204 Pb-206 Pb/ 204 Pb closure, the uranogenic Pb isotope compositions (e.g., 206 Pb/ 204 Pb) of the Kuetsjärvi lavas were markedly unradiogenic relative to contemporaneous average crustal and mantle reservoirs, yet the thorogenic Pb ( 208 Pb/ 204 Pb) was similar to coeval mantle (Fig. 2a-d; (43)). These relations indicate that between eruption ~2.06 Ga and metamorphism ~1.80 Ga, 208 Pb/ 204 Pb increased via decay of 232 Th while 206 Pb/ 204 Pb and 207 Pb/ 204 Pb remained invariant (Fig. 2c). This requires near-complete to complete removal of U, resulting in little to no accumulation of uranogenic Pb (Fig. 2d), concurrent with the uninterrupted accrual of thorogenic Pb in these rocks. The Pb isotope systematics therefore document early, selective removal of U via oxidation while Th remained immobile. This is the only reasonable explanation for why the KVF 208 Pb/ 204 Pb-206 Pb/ 204 Pb array plots above the S-K curve while the KVF 207 Pb/ 204 Pb-206 Pb/ 204 Pb array plots below the S-K curve (Fig. 1), in the absence of evidence for addition of Pb (above).
The removal of U from Unit E is likely tied to chemical weathering at the surface and/or alteration in contact with oxidized groundwaters. This is consistent with the pervasive secondary oxidation documented throughout the Kuetsjärvi Formation (28,(49)(50)(51), which was interpreted to immediately postdate eruption based on incorporation of oxidized lava clasts in the overlying Kolosjoki Sedimentary Formation (maximum age 2058 ± 6 Ma; (52)). Crucially, soluble U 6+ would sorb or coprecipitate with any Fe 3+ -oxide/hydroxide phases that may have been present (e.g., (61,62)), requiring that the primary U in these basalts was oxidized and leached away before secondary Fe 3+ -minerals formed during oxidative weathering and alteration (SI Appendix). To quantitatively assess the relative rates of U and Fe oxidation in the KVF, a simple reaction path model was developed (Fig. S16, Table S6). The simulation results support the hypothesis that microcrystalline uraninite or glass-bound U would undergo oxidative dissolution before significant Fe 2+ -silicate oxidation and associated ferrihydrite formation. In contrast, Th concentrations would have remained comparably unchanged during oxidative weathering and alteration, apart from the allanite-bearing sample at 97.52 m.
While there was minimal to no ingrowth of uranogenic Pb in the KVF between 2.06 and 1.80 Ga (Fig. 2d), the 207 Pb/ 204 Pb-206 Pb/ 204 Pb isochron age indicates that ingrowth of uranogenic Pb resumed at ~1.80 Ga. The initial 206 Pb/ 204 Pb and 207 Pb/ 204 Pb ratios when uranogenic Pb ingrowth resumed are estimated to lie slightly below that of the mantle at 2.06 Ga (Fig. 2b), reflecting a "frozen" isotopic composition after removal of U shortly after eruption at 2.06 Ga. Reintroduction of U is most likely to have occurred during the ~1.8 Ga Svecofennian Orogeny, where oxidized shallow U 6+ -bearing fluids would have been reduced upon encountering greenschist-facies metamorphic conditions (SI Appendix, Fig. S17). Re-introduction of U was variable, producing a range in measured Th/U ratios (3.2-5.8) that significantly exceeded the relatively narrow range expected for mantle-derived basalts, and this is confirmed by the wide range in measured U/Zr ratios (SI Appendix, Fig. S10). The distributions of U, Th, Pb, and Fe across the leaches (SI Appendix, Fig. S12) suggest that the added U was nearly equally distributed between recrystallized silicates/accessory minerals (represented by the HF leaches) and Fe 3+ -bearing oxides (as represented by the HCl leaches), consistent with metamorphic U addition to greenschist-facies rocks that contained high abundances of hematite and magnetite.
It perhaps comes as no surprise that the Kuetsjärvi lavas have experienced a complex history given the tectonothermal history of the region. Despite this, a full U-Th-Pb isotope analysis provides multiple constraints in Pb-Pb space and four sets of age information, that, in addition to elemental data, allows us to decipher the timing of redox changes in U as related to post-eruption weathering/alteration and oxidation despite subsequent metamorphic re-introduction of U. We infer that oxidative removal of all, or nearly all primary U at ~2.06 Ga from the KVF basalts was a process not isolated to this locality, and instead represents general, global conditions. The period between ~2.2 Ga and ~1.8 Ga represented a strong increase in subaerial exposure of basaltic rocks, culminating in establishment of present-day continental freeboard (63), suggesting that deep oxidative weathering and alteration of basalts at ~2 Ga would have produced a large dissolved U flux to the oceans. It is also possible that oxidative weathering of granitic-composition crust may have released U 6+ to the oceans during the GOE, although there are some difficulties in assessing this possibility, most important of which is that much of this U may have been locked up in accessory minerals such as zircons that would have been resistant to oxidative weathering. Regardless of the sources of U 6+ to the oceans during the GOE, in the next section we speculate on the relevance of this intense chemical weathering period and its long-lived consequences for the surface and deep Earth in terms of the U-Th-Pb isotope systematics of modeled mantle reservoirs (Figs. 3, 4a) and global sedimentary U geochemistry (Fig. 4b-e).
A comparison of the U-Th-Pb isotope systematics of oxidatively weathered/altered Kuetsjärvi lavas with typical mantle reservoirs is most simply visualized by quantifying the offset in 207 Pb/ 204 Pb and 208 Pb/ 204 Pb relative to 206 Pb/ 204 Pb. Figure 3 shows Δ7/4 and Δ8/4 Pb isotope anomalies of the Kuetsjärvi lavas relative to the Northern Hemisphere Reference Line (NRHL; (64)). Modern mid-ocean ridge basalts (MORBs) and ocean-island basalts (OIBs) exhibit a lowgradient Δ7/4-Δ8/4 array, consistent with heterogeneity in the Th/U and age of their sources. In contrast, the Kuetsjärvi Unit E lavas have exceptionally high positive Δ8/4 values relative to the NHRL and the MORB-OIB Δ7/4 -Δ8/4 array, which can only be explained by decoupling of the U-Pb and Th-Pb isotope systems in the Kuetsjärvi rocks. The very high Δ8/4 values (142 to 319, with an outlier at 1108) reflect ingrowth of thorogenic but not uranogenic Pb from 2.06-1.80 Ga (Figs. 2c, 2d, 3), consistent with the model of oxidative U loss in the rocks after eruption. The relatively low Δ7/4 values (-17 to -12) confirm the lack of accrual in uranogenic Pb in these rocks from 2.06 to 1.80 Ga while the evolving mantle reference frame continuously accrued uranogenic Pb. Importantly, the contrast in Δ7/4 and Δ8/4 persists through re-introduction of U during metamorphism at 1.8 Ga, where the extent of U addition would control the absolute values of Δ7/4 but would not significantly change the relative Δ7/4 and Δ8/4 values. Late Phanerozoic Pb loss had no meaningful effect on the relative Δ7/4 and Δ8/4 values.
Our proposed weathering flux of U 6+ to the oceans, whose loss from the KVF has produced positive Δ8/4 values, appears to be mass-balanced by the negative Δ8/4 values of the HIMU (high-238 U/ 204 Pb, or "μ") mantle reservoir (Fig. 3). Although relative enrichment of U in the HIMU reservoir is fundamental to its definition (e.g., 38,65), our results identify a potential massbalance for U-loss through oxidative weathering on the continents. These observations support geochemical evidence and modeling that suggest incorporation of excess U relative to Th in the mantle ~2 Ga, including solutions to the kappa (κ; 232 Th/ 238 U) conundrum, which identifies a mismatch in the measured and modeled κ of the MORB source (39,66), and the existence of the HIMU reservoir contributing to OIBs (67,68).
The 'kappa conundrum' for MORBs reflects the fact that the average measured κ of MORBs (κ~2.5) is lower than the time-integrated model κ of the MORB source region required by 208 Pb/ 204 Pb-206 Pb/ 204 Pb systematics (κ~3.8). A potential resolution is gradual enrichment of U relative to Th in the MORB mantle source via subduction of altered oceanic crust (AOC; (69)). In Figure 4a, the family of solutions for evolution of κ in the MORB mantle that satisfy the κ conundrum is shown, with a shift from κ ~4.0 before 2.5 Ga (Fig. 4a; (69)) to κ ~2.5 today. The inflection to lower κ after 2.5 Ga corresponds with rising atmospheric oxygen levels (subsequent studies model this at 2.35 Ga; e.g., (70,71)), which is consistent with the onset of oxidative weathering and mobilization of soluble U from the crust to the oceans. Ultimately, subduction of AOC after the GOE would introduce low-κ, high-μ material into the mantle. Our results for the Kuetsjärvi lavas document, for the first time, direct evidence for release of soluble U during oxidative weathering/alteration of continental crust, consistent with models suggesting enrichment of U in AOC during the early Paleoproterozoic (70)(71)(72), as well as models of κ evolution through time (69,73).
The high 206 Pb/ 204 Pb ratios that characterize HIMU and some OIBs, relative to MORBs, have long been taken as evidence for enrichment of U relative to Th in the mantle ~2 Ga (e.g., (74,75)). Mechanisms to produce U enrichment in the sources of HIMU are varied, and include: direct remelting of subducted altered oceanic crust (AOC) with high μ (65,(75)(76)(77); reprocessing of AOC as it is subducted, with resultant metasomatism of mantle domains which are subsequently remelted (71); delamination or remelting of metasomatized subcontinental lithospheric mantle (78,79); or involvement of a carbonate-rich component (80). A number of workers have focused on intra-mantle processes that may modify AOC as it is subducted, pointing to depletion in LILEs as an indicator of dehydration of AOC during subduction and positive Nb-Ta anomalies as an indicator of residual subducted oceanic crust (65), as well as Sr isotope geochemistry (e.g., (68,73)).
We suggest that formation of HIMU mantle sources at ~2 Ga directly correlates with an increased flux of soluble U to the oceans via contemporaneous oxidative weathering/alteration of the continents, making a connection between oxidative weathering of exposed continental crust and the geochemistry of Paleoproterozoic AOC. Multiple processes govern the relative fractionation of U, Th, and Pb in AOC as it is progressively subducted into the mantle. The fluids released from AOC to arcs have low U/Pb ratios, consistent with high-U/Pb residual material (e.g., (81)). While the deep oceans were likely not fully ventilated until the Neoproterozoic Oxygenation Event (e.g., (82)), uptake of U into ocean crust could have been localized proximal to the continents, and it has been shown in modern systems that slab bending and tearing at subduction zones creates channels and pathways for water to hydrate the crust (83). A possible example of U-enriched oceanic crust may be found in 1.9 Ga altered oceanic basalts of the Flin Flon assemblage, Canada, where high 206 Pb/ 204 Pb relative to 208 Pb/ 204 Pb has been interpreted as evidence of enrichment of U during alteration of the ocean crust (72,84). While this provides a clear instance of U enrichment into Paleoproterozoic oceanic crust, the mechanisms for its transport into the mantle are not well constrained.
Juxtaposition of the ~2.06 Ga oxidative weathering/alteration of the Kuetsjärvi lavas and changes in mantle Th/U (Fig. 4a) with broader indicators of redox changes in the atmosphere and marine systems (Figs. 4b-d) demonstrates potential interconnections between the surface and deep Earth through the LJE. Net organic carbon burial on the continents may be used as a proxy for atmospheric O2 levels (85), and predict rising atmospheric O2 from ~1% PAL at 2.5 Ga to ~4% PAL at 2.06 Ga, when the Kuetsjärvi lavas were emplaced and oxidatively weathered/altered (Fig. 4b). This fourfold increase in atmospheric O2 is coincident with geochemical changes in shales, including elevated [U] (Fig. 4c; (17,18)) and δ 238 U values (Fig. 4d: (18,(86)(87)(88)(89) at ~2.1-2.0 Ga. In marine systems, input of soluble U 6+ via hydrothermal systems is minimal, so the reduction and uptake of U into anoxic sediments tracks release of U via oxidative weathering on the continents (e.g., 82,83). High-U organic-rich shales from this interval (Fig. 4c; (17,18)) include the ~2.1-2.0 Ga organic-rich Zaonega Formation of the Onega Basin, Russia (up to 238 ppm U; (18,86), and the Francevillian Group, Gabon (up to 75 ppm post-depositional enrichment of U in unit FA and 8 ppm primary U enrichment in unit FD; (92,93)), suggesting that the oxidative weathering/alteration recorded in the Kuetsjärvi volcanics may correspond with a global, substantial riverine flux of oxidized U to the oceans at this time. Formation of the extraordinary Oklo natural nuclear fission reactors of the Francevillian Basin, Gabon at 2050 ± 30 Ma (Fig. 4e; (94)) is also testament to the presence of oxidized groundwater ~2.1-2.0 Ga that transported U 6+ , followed by reductive precipitation and concentration in organic-rich rocks (95,96).
Uranium enrichment in Paleoproterozoic organic-rich shales is most pronounced for the ~2.06 Ga Zaonega Formation of the Onega Basin, Russia, whose depositional age overlaps that of the Kuetsjärvi lavas. The Onega Basin is ~800 km SW of the Pechenga Belt, and the high TOC content of these rocks (up to 70 wt%; 5) has resulted in elevated authigenic U and a wide range of δ 238 U values (Fig. 4c,d), interpreted to reflect decreasing atmospheric O2 (86) or high atmospheric O2 in the wake of the Lomagundi-Jatuli Event (18). Worldwide, the shale and carbonate records track minimal U isotope fractionation from 3.0 to 2.0 Ga (97), which may be taken to indicate either a relatively minor flux of oxidized U to the oceans or quantitative uptake of U upon reductive sequestration. It is difficult to independently estimate the relative fluxes of U involved in oxidative and reductive pathways on the surface of the Earth during this time. The range of δ 238 Uauth preserved in Zaonega Formation shales seems likely to, in part, indicate high atmospheric oxygen conditions during its deposition (18), although interpretations may be complicated given the apparent disconnect with the carbonate record (92,97), pending refinement of age constraints for carbonates in the Onega Basin (98). Our results suggest that soluble U 6+ was abundant, and that the contrasting records in shale and carbonate δ 238 U values may indicate variable reductive U sequestration in marine shales.
The distinctive low-207 Pb/ 204 Pb and -206 Pb/ 204 Pb, high-208 Pb/ 204 Pb ratios of the basaltic Kuetsjärvi Unit E lavas (Figs. 1, 3) document an excess of thorogenic 208 Pb that can only be explained by early suppressed ingrowth of uranogenic 207 Pb and 206 Pb. The Pb isotope compositions require early Th/U disturbance in the rocks via complete or near-complete U loss during oxidative weathering/alteration shortly after eruption at ~2.06 Ga. Comparison of 207 Pb-206 Pb, 238 U-206 Pb, 235 U-207 Pb, and 232 Th-208 Pb ages indicate that re-introduction of U occurred during metamorphism at ~1.8 Ga (Fig. 2), followed by Pb loss between ~400 and 200 Ma. Uranium loss during oxidative weathering/alteration of the basalts at ~2.06 Ga corresponds with significant authigenic enrichment of U in marine sediments on a broad scale (Fig. 4), which approximately corresponds with geochemical evidence and modeling that suggest the incorporation of excess U relative to Th in the mantle ~2 Ga. Our results therefore point toward a potential mechanism for the change at ~ 2 Ga in κ of the MORB source (Fig. 4) and the existence of the HIMU reservoir (Figs. 3, 4), implying a significant connection between the redox state of surface-and deep-Earth reservoirs in the Paleoproterozoic. Our U-Th-Pb isotope results support proposals of elevated atmospheric oxygen concentrations during the Lomagundi-Jatuli Event. Moreover, we show how detailed geochronology may document mobility of redox-sensitive trace elements over multiple geological events. In no case do the measured U abundances in the samples capture this complexity, indicating that by themselves, elemental abundances may not provide clear proxy information in rocks that have undergone complex tectonothermal histories. of mantle reservoirs. Gray region is range of models that solve the kappa conundrum for MORB mantle evolution (69). Black lines are models of high-μ (HIMU) and depleted mantle (DM) evolution, with the HIMU reservoir becoming isolated from DM ~2 Ga (73). B) Global organic carbon burial model that accounts for sedimentary volume flux and of distributions of total sedimentary organic carbon content by lithology type, with the median variant model using the variable North American record as a proxy for global organic carbon burial (85). Predicted atmospheric O2 levels of 1% and 10% PAL (Present Atmosphere Level) are from their model of sinks and sources of atmospheric oxygen. C) [U] in shales (17,18). Dark gray symbols are from the proximal Onega Basin, Russia at 2.06 Ga (101,102). D) δ 238 U in shales (18,(86)(87)(88)(89). Lightest gray symbols are ~1.73 Ga Wollogorang Fm. shales which may have been hydrothermally overprinted (89). E) Timeline of key tectonic and climate events (19, 52-54, 94, 103, 104). 208 Pb/ 204 Pb 207 Pb/ 204 Pb 206 Pb/ 204 Pb 238 U/ 204 Pb 2 S.E. 2σ 1 232 Th/ 204 Pb 2 S.E. 2σ 1 Core Sample # (m) unconformity (m) (wt%) (ppm) (ppm) (ppm) (%) (%) 6A 3114734 97.36 0 13.18 0.69 5.27 0.82 1.30 40.383 15.321 17.864 10.14 0.01 0.07 16.50 0.21 0.24 6A 3114736 97.52 0.16 15.02 0.58 8.20 2.82 24.28 52.435 15.666 21.281 27.18 0.02 0.07 239.62 0.28 0.31 6A 3114738 97.58 0.22 16.06 0.58 7.13 0.96 3.90 38.910 15.246 17.203 8.50 0.20 0.21 35.46 0.09 0.15 6A 3114740-2 97.74 0.38 20.17 0.63 5.00 0.74 4.21 40.058 15.260 17.204 9.56 0.02 0.07 55.52 0.09 0.15 6A 3114742-1 99.08 1.72 18.03 0.59 4.75 0.79 3.15 38.986 15.313 17.844 10.63 0.01 0.07 43.45 0.08 0.15 6A 3114742-2 99.13 1.77 22.28 0.69 5.58 0.87 3.40 39.120 15.315 17.791 9.98 0.01 0.07 40.01 0.09 0.15 6A 3114744-1 101.04 3.68 15.33 0.60 3.41 0.95 2.96 40.542 15.490 19.386 18.62 0.04 0.08 59.45 0.08 0.14 6A 3114744-2 101.09 3.73 14.98 0.57 4.60 0.95 3.03 41.365 15.536 19.676 13.94 0.02 0.07 45.84 0.06 0.14 6A 3114744-3 101.14 3.78 15.87 0.59 4.85 0.63 2.84 39.551 15.349 18.011 8.36 0.02 0.07 38.78 0.07 0.14 1 Two-sigma uncertainties for 238 U/ 204 Pb and 232 Th/ 204 Pb include propagation of excess 0.07% and 0.12% uncertainty, respectively.