A review of temporal constraints for the Palaeoproterozoic large, positive carbonate carbon isotope excursion (the Lomagundi–Jatuli Event)

Earth-Science Reviews 127 (2013) 242–261

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A review of temporal constraints for the Palaeoproterozoic large, positive carbonate carbon isotope excursion (the Lomagundi–Jatuli Event) Adam P. Martin a,⁎, Daniel J. Condon a, Anthony R. Prave b, Aivo Lepland c,d a

NERC Isotope Geosciences Laboratory, British Geological Survey, Keyworth, Nottinghamshire, NG12 5GG, UK Department of Earth and Environmental Sciences, University of St Andrews, St Andrews, KY16 9LF, UK Geological Survey of Norway, Postboks 6315 Slupen, 7491 Trondheim, Norway d Tallinn Technical University, Institute of Geology, 19086 Tallinn, Estonia b c

a r t i c l e

i n f o

Article history: Received 4 March 2013 Accepted 9 October 2013 Available online 18 October 2013 Keywords: Paleoproterozoic U–Pb geochronology Lomagundi–Jatuli Event Carbon isotope excursion Great Oxidation Event

a b s t r a c t The Palaeoproterozoic Lomagundi–Jatuli Event is one of the largest magnitude and earliest known positive carbonate carbon isotope excursions, preserving δ13C values between +5 and +16‰ and even higher. It is recorded in sedimentary rocks on all continents bar Antarctica and spans stratigraphic thicknesses ranging from several to many tens of metres. This unique positive δ13C interval signals fundamental changes in the global carbon cycle and is a key event in Earth system evolution following oxygenation of Earth's atmosphere. Here we present a comprehensive review of the age constraints on the Lomagundi–Jatuli Event, the first such effort in two decades. This new chronology compilation focuses on the U–Pb and Re–Os chronometers and demonstrates that global synchronicity of the Lomagundi–Jatuli Event is permissible, an interpretation supported by the apparent wide dispersion of Lomagundi–Jatuli Event-bearing successions in the most recent Palaeoproterozoic plate reconstructions. Assuming the Lomagundi–Jatuli Event is synchronous worldwide, then the bounds on its duration range from a maximum of 249 ± 9 Myr (2306 ± 9 Ma to 2057 ± 1 Ma) to a minimum of 128 ± 9.4 Myr (2221 ± 5 Ma to 2106 ± 8 Ma). © 2013 Elsevier B.V. All rights reserved.

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Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Background on the Lomagundi–Jatuli Event . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Oxygen and the Great Oxidation Event . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2. The paradox of 13C-enriched carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3. Hypotheses on the origins of the Lomagundi–Jatuli Event . . . . . . . . . . . . . . . . . . . Palaeoproterozoic tectonic reconstruction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochronometers and approaches appropriate for dating Precambrian stratigraphy . . . . . . . . . . . 4.1. Challenges in Precambrian geochronology . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2. Suitable geochronometers and dateable material . . . . . . . . . . . . . . . . . . . . . . . 4.3. A review of maximum, minimum and direct age constraints . . . . . . . . . . . . . . . . . . Chronological constraints on the Lomagundi–Jatuli Event . . . . . . . . . . . . . . . . . . . . . . . 5.1. North America . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2. Russia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3. Finland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.4. Africa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5. Australia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.6. South America . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.7. Scotland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.8. Other localities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1. Duration of the Lomagundi–Jatuli Event . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2. Global versus diachronous nature of the Lomagundi–Jatuli Event . . . . . . . . . . . . . . . . 6.3. Timing of the Lomagundi–Jatuli Event compared to other key Archaean to Palaeoproterozoic events

⁎ Corresponding author at: GNS Science, Private Bag 1930, Dunedin, New Zealand. E-mail address: [email protected] (A.P. Martin). 0012-8252/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.earscirev.2013.10.006

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7. Targets for future chronological work 8. Conclusions and concluding remarks . Acknowledgements . . . . . . . . . . . References . . . . . . . . . . . . . . .

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1. Introduction In the late 1960's and mid 1970's Precambrian sedimentary carbonate rocks were measured systematically for their carbon isotopic compositions from the Lomagundi Group in Zimbabwe and the Jatuli ‘group’ in Fennoscandia (Galimov et al., 1968; Schidlowski et al., 1975). These studies revealed surprisingly high δ13C values, as high as +13‰ (all values of δ13C reported here refer to carbonate carbon normalised to the Vienna Pee Dee Belemnite standard). Initially the African and Fennoscandian values were considered to be local phenomena (Galimov et al., 1968; Schidlowski et al., 1976) and it was not until similarly high values were documented in Scotland and Norway (Baker and Fallick, 1989a,b) that they were thought to record a global event, an event later named as the Lomagundi–Jatuli positive carbonate carbon isotope excursion (Melezhik et al., 2005) or Lomagundi–Jatuli Event (LJE). A compilation of marine carbonate isotope data across Earth history (Shields and Veizer, 2002; Fig. 1) reveals how anomalous the LJE values are, commonly N+5‰ to +16‰ (Melezhik et al., 1997) and even as high as +28‰ (Bekker et al., 2003a), when compared to the typical range in global δ13C of between −5‰ and +5‰. The first attempt to establish the duration of the LJE was provided by Karhu (1993) and Karhu and Holland (1996) using a compilation of absolute ages from 22 sections in Fennoscandia, Norway, Scotland, Africa, North America and Australia. From those they determined a time span of 2220 Ma to 2060 Ma for the LJE and a definition that has become widely adopted for the LJE is the presence of a primary, positive δ13C signal N +5‰ in carbonate rocks deposited during the Siderian to Rhyacian Periods. Since the Karhu and Holland (1996) publication a significant number of radiometric ages, utilising enhanced methods with improved accuracy and precision, have been obtained on many LJE-bearing sections worldwide making an updated review of the geochronology of the LJE at this time appropriate. In this paper a comprehensive review of the temporal constraints on stratigraphic successions with δ13C data is presented and discussed in the context of the LJE. A summary of relevant isotopic data, hypotheses of the genesis of the LJE, Archaean and Palaeoproterozoic plate configurations and overview of appropriate Precambrian chronometers and their use as constraints to the age of sedimentary sections are also discussed. The radiometric ages are considered both by region and as a global compilation, and are used to assess the timing of the initiation, duration, termination and potential global synchronicity of the LJE. Lastly, future directions for chronological research are highlighted.

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(Cloud, 1968), as based on the appearance of oxidised ‘red’ beds in sedimentary successions and the disappearance of detrital minerals such as uraninite, pyrite and siderite that are unstable under oxygenated conditions (Cloud, 1972 and references therein). The geological data have been complemented by geochemical evidence for localised increases of oxygen in the surficial environment as early as 2.5 Ga from peaks in concentration of redox sensitive elements such as Mo and Re in black shales (Anbar et al., 2007; Wille et al., 2007) and Cr in pyrite (Konhauser et al., 2011). However, the marked oxygenation of Earth's atmosphere (the Great Oxidation Event; Holland, 2002) to levels of O2 above 10−5 of present atmospheric levels (Pavlov and Kasting, 2002) is generally defined as the end of mass independent fractionation of S isotopes (Farquhar et al., 2000; Farquhar and Wing, 2003). This is constrained to have occurred between c. 2.45 and 2.32 Ga (Bekker et al., 2004; Guo et al., 2009; Fig. 1). The LJE follows on from the Great Oxidation Event, but as yet its genesis remains enigmatic and exact timing only constrained broadly.

2. Background on the Lomagundi–Jatuli Event The LJE is one of a suite of hallmark events of the Archaean– Proterozoic transition (Fig. 1) that includes the end of massindependent fractionation of sulphur isotopes and the global rise of free oxygen, deposition of the earliest non-marine ‘red beds’ and repeated worldwide glaciations (Kasting, 1993; Aharon, 2005; Canfield, 2005; Kump, 2008; Pufahl and Hiatt, 2012; Melezhik et al., 2013 for a recent review). The first-order temporal association of these events has resulted in a number of hypotheses that seek to explain the sustained positive δ13C values of the LJE in the context of the rise of atmospheric oxygen and major environmental change. 2.1. Oxygen and the Great Oxidation Event The Archaean–Palaeoproterozoic transition has long been proposed as a time when free atmospheric di-oxygen became relatively abundant

Fig. 1. Schematic depiction of major events during the Archaean to Proterozoic transition. The dark and light boxes in the upper panels represent reducing and oxidising atmospheres, respectively; PAL = present atmospheric levels; S = sulphur (after Holland, 1994; Bekker and Holland, 2012). Recent data indicate a transient rise in oxygen levels (of 0.04 to 0.1 pO2) between 2.45 and 2.2 Ga before present. The distribution of Fe deposits is after Isley and Abbott (1999). Plot of Δ33S versus age after Farquhar and Wing (2003) and Reinhard et al. (2013); white circles are bulk rock data and grey circles are secondary ion mass spectrometry generated data. Carbon isotopic evolution of marine carbonates is after Shields and Veizer (2002).

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2.2. The paradox of 13C-enriched carbonates Firm knowledge of the links between organic carbon production and burial, carbonate deposition and redox conditions is crucial in understanding the carbon isotope record and its use as a proxy in understanding the evolution of atmospheric (and, hence, oceanic) composition. Organic matter produced by photosynthesis is significantly enriched in the lighter isotope 12C, leaving the (inorganic) carbonate reservoir enriched in 13C (O'Leary, 1981). The pattern of C-isotopic values from 3.5 Ga onwards is remarkably consistent, with δ13C values of carbonate fluctuating around an average of 0 ± 5‰ and δ13C values of organic matter at ~−26 ± 7‰ (Shields and Veizer, 2002; Fig. 1). This suggests that the organic and inorganic carbon reservoirs have been deposited in a ratio of about 1:4 since at least 3.5 Ga (Aharon, 2005), with 20% of the total carbon pool as organic carbon (Shaw, 2008). The LJE represents a major disturbance to this system, which has been hypothesised as reflecting a coeval increase in the overall size of the organic carbon pool (Karhu and Holland, 1996). However, the LJE occurs commonly in successions in which there is a paucity of contemporaneous organic-rich sediments preserved or reported (Karhu and Holland, 1996; Buick et al., 1998; Melezhik et al., 1999; Bekker et al., 2001; Lindsay and Brasier, 2002), the so-called Palaeoproterozoic ‘paradox’ of Melezhik et al. (1999). 2.3. Hypotheses on the origins of the Lomagundi–Jatuli Event Mechanisms to form and preserve high δ13C carbonates can be ascribed to a combination of tectonic and biospheric processes (Shaw, 2008). Many workers appeal to variations in organic carbon burial rates, fluctuations in ocean chemistry (redox gradients), changes in metabolic pathways and/or diagenetic methanogenesis response to the oxygenation of the Earth, all of which can be affected by continental assembly and rifting as these have a first order impact on geochemical cycling (Yudovich et al., 1991; Dix et al., 1995; Karhu and Holland, 1996; Shields, 1997; Aharon, 2005; Hayes and Waldbauer, 2006; Bekker et al., 2008; Kirschvink et al., 2009). Karhu and Holland (1996) and Bekker et al. (2008) have favoured high organic carbon burial rates that caused global shifts in isotopic composition of the atmospheric CO2 to higher δ13C values. Aharon (2005) calculated that an organic burial flux three times greater than in the Phanerozoic would be required to account for the high LJE values. However, as noted above, organic-rich units correlative with LJE-bearing successions are scarce. Other speculative scenarios include envisioning enhanced nutrient fluxes (P) caused by more acidic weathering under Palaeoproterozoic atmospheric conditions to generate a global increase in productivity (e.g. Konhauser et al., 2011; Bekker and Holland, 2012). A creative alternative model has been proposed by Schrag et al. (2013) who argue that authigenic carbonate, produced during early diagenesis, is a largely overlooked major component of the carbon mass balance equation, particularly for Precambrian strata. There is also a debate on whether the LJE is a singular, global excursion or one comprised of several isotope excursions and even if it represents a number of regionally restricted and diachronous events (Karhu and Holland, 1996; Buick et al., 1998; Melezhik et al., 1999; Bekker et al., 2006; Bekker and Kaufman, 2007; Melezhik et al., 2007; Maheshwari et al., 2010). An objective of this review is to synthesise and assess the geochronological constraints on the LJE to help better delineate the architecture of the LJE and differentiate between the various hypotheses. 3. Palaeoproterozoic tectonic reconstruction The supercontinent cycle links processes from Earth's surface to deep interior and is a key component of several hypotheses attempting to explain the driving force behind the LJE. Palaeomagnetic data and peaks in the compilations of zircon age distributions have been commonly used as evidence for the existence of Proterozoic supercontinents

(e.g. Condie, 1998; Evans and Powell, 2000; Evans et al., 2001; Evans, 2002; Hawkesworth et al., 2010) such as Kenorland (Williams et al., 1991; Aspler and Chiarenzelli, 1998; Bleeker, 2003), Vaalbara (Cheney, 1996; Bleeker, 2003) and Arctica (Rogers and Santosh, 2003) among others. However, stratigraphic and palaeomagnetic viability of these purported single, early supercontinents, particularly Kenorland, has come under increasing scrutiny and questioning, with numerous workers suggesting instead a wide dispersal of cratons between 2500 and 2000 Ma (e.g. Evans, 2009; Reddy and Evans, 2009; de Kock et al., 2009; Eglington et al., 2013; Pehrsson et al., 2013). The most recent plate reconstruction for the Palaeoproterozoic, in this case at 2100 Ma, is shown in Fig. 2 (Eglington et al., 2012, 2013). 4. Geochronometers and approaches appropriate for dating Precambrian stratigraphy Establishing chronologies in stratigraphic successions that contain rocks suitable for radio-isotopic dating (i.e. interstratified felsic volcanic rocks) can be relatively straightforward. However, such occurrences with well established relationships to LJE-bearing and associated carbonate successions are few. This necessitates using a variety of approaches that require circumspection regarding how they are integrated. In this section we briefly outline the types of radio-isotopic data used to date Precambrian stratigraphic successions in general, and outline the challenges for constraining the timing of the LJE in particular. 4.1. Challenges in Precambrian geochronology Stratigraphic timescales for Phanerozoic successions are developed based upon a combination of highly resolved bio-, chemo- and magneto-stratigraphic records that are integrated regionally and globally correlated and constrained in an absolute sense by radioisotopic dates (primarily mineral ages from inter-stratified ash beds). In contrast, most of the Precambrian stratigraphic record lacks bioand magneto-stratigraphies and is reliant solely upon radio-isotopic dating to establish absolute age, and to develop temporal frameworks within which data from disparate successions can be integrated. Whilst there is potential for some correlation based upon chemo- and lithostratigraphy (e.g. glacial deposits, δ13C excursions, etc.) many questions remain about synchronicity. Further, using them for the purpose of chronostratigraphic correlation introduces an element of circular reasoning. Further, the c. 2.4–2.2 Ga time slice is the least represented in terms of crustal growth in the geological record (see Hawkesworth et al., 2010 and references therein for a discussion). Thus, not only are the methods for building a temporal framework restricted to radio-isotopic ones, the preserved geological material to which they can be applied is also scarce. 4.2. Suitable geochronometers and dateable material Condon and Bowring (2011) review extensively approaches for constraining Precambrian (and Phanerozoic) stratigraphy and their salient points for materials and radio-isotopic systems suitable for constraining Precambrian successions are summarised here: a) The U–Pb decay scheme applied to U-bearing accessory minerals from volcanic and/or intrusive igneous rocks (zircon, titanite, baddeleyite, etc.) where the dated mineral is inferred to date the crystallisation age of the rock. b) The U–Pb decay scheme applied to detrital U-bearing minerals (e.g. zircon) where the dated mineral is older than the rock in which it was contained and provides information about provenance and crustal evolution as well as a maximum depositional age constraint for the level sampled. c) The U–Pb or Re–Os decay schemes applied to chemical precipitates and organic residues such as carbonates, phosphates and/or

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Fig. 2. 2100 Ma plate reconstruction (Eglington et al., 2012, 2013) with the position of relevant carbonate sections marked by a star symbol. 1. Wyoming, 2. Huronian, 3. Labrador Trough, 4. Scotland, 5. Zimbabwe, 6. Transvaal, 7. Western Australia, 8. India, 9. China, 10. Rio de la Plata, 11. Gabon, 12. Sao Francisco, 13. Norway, 14–19 Fennoscandia.

organic-rich sediments. In general this approach requires enrichment in a parent element (e.g. U, Re) and spread in the ratio of parent to daughter isotope so an isochron can be constructed. An exception is the U–Pb dating of diagenetic xenotime overgrowths on detrital zircons (Rasmussen, 2005), which does not require construction of an isochron. All these approaches have their own set of assumption and caveats. Key, though, is that the material dated is a robust chronometer, has remained a closed system since formation and there has been no mobility of parent and/or daughter isotopes. For the U–Pb system, the concordance of the two decay schemes can be used to quantify the degree to which the system has remained closed. In practical terms this is complicated by the uncertainty of the different U–Pb dates and a particular issue for Palaeoproterozoic (and older) U–Pb data are non-zero Pb-loss. For the isochron methods a statistical measure of the fit of the data to a line (i.e. isochron) is often used to assess ‘goodness’ but this also has a variety of uncertainties. The approach in this study has been to focus on data from U–Pb and Re–Os studies and to exclude Pb–Pb, Rb–Sr or Sm–Nd whole-rock ages where closed system behaviour is difficult to assess independently. In that the latter methods typical have large uncertainties (many tens of millions of years), their exclusion has little impact on the assessment of the temporal constraints of the LJE and are of limited utility compared to the more precise U–Pb and Re–Os ages. Similarly, we have not included systematic uncertainties related to constants used in the calculation of dates. These would include the 238U, 235U and 187Re decay constants and the natural 238U/235U ratio. In this compilation 207Pb/206Pb dates are relative to a 238U/235U ratio equal to 137.88 and using the 238U and 235U decay constants of Jaffey et al. (1971). Recent work (Hiess et al., 2012) indicates a 238U/235U value of 137.818 ± 0.045 is more appropriate for zircon U–Pb (and 207Pb/206Pb) geochronology and this value has the effect of lowering all Palaeoproterozoic 207Pb/206Pb dates by c. 0.8 ± 0.6 Myr, but is not significant given the magnitude of other sources of uncertainty. The Re–Os uncertainties in this compilation (Puchtel et al., 1999; Hannah et al., 2004) do include uncertainty in the 187Re decay constant estimated at 0.31%. Again, these sources of systematic uncertainty are not significant given that the analytical and geologic sources of uncertainty are about an order of magnitude greater. 4.3. A review of maximum, minimum and direct age constraints Once a radio–isotopic age constraint has been obtained for a given mineral or rock sample the next question is how does it relate to the geological feature of interest, in this case the δ13C profile of carbonate

rocks? Chronological constraints for LJE-bearing sedimentary strata rely upon the presence of dateable materials (tuffs, lava flows, igneous bodies with peperitic margins, authigenic/diagenetic minerals) and how those relate to the LJE rocks. Clearly intercalated volcanic rocks would be relatively straightforward to interpret as direct age constraints on the timing of the LJE whereas other types of data from detrital minerals and/or intrusive rocks would provide information about minimum and maximum age constraints. Considering a hypothetical stratigraphy (Fig. 3) in which a LJEbearing unit (B) is under- and over-lain by sedimentary units (A, C) that are marked by isotopically normal δ13C values and all three units contain an extrusive and/or peperitic igneous body, an intrusive (cross-cutting) igneous body and a detrital zircon-bearing interval. How the temporal information contained within each unit (maximum, minimum and/or direct ages) constrains the interval of interest depends on the stratigraphic relationships. For example (Fig. 3), a date from the youngest detrital zircon grain in unit B presents a maximum age constraint for the LJE, a date from an intercalated tuff or lava would present a syn-depositional age, whereas a date from a cross-cutting igneous rock provides a minimum age constraint. However, for unit C a detrital mineral date provides no constraint on the age of unit B, a date from an extrusive or cross-cutting igneous body unit provides a minimum age, and so forth. As highlighted by Fig. 3, a purely objective assessment of obtained ages requires treating the stratigraphic position of a dated sample in relation to the sampled interval of interest without (additional) inference of the geological relationships between each unit. Beyond this, it is necessary to consider the geological context. For example, applying the above criteria exactingly, in the case where an igneous rock cross-cuts a siliciclastic formation (unit A) that grades upward transitionally into carbonate rocks containing the LJE (such as Fig. 3 unit B), the intrusion would not constrain the age of the carbonate strata. However, if it can be reasonably shown that deposition between units A and B was continuous (without evidence for a significant hiatus) then it can be inferred that the intrusion dated in unit A also provides a minimum age constraint to the LJE-bearing rocks. Such reasoning has been used to constrain the timing of the LJE on Fennoscandia (Melezhik et al., 2007; Martin et al., 2013): detrital zircons obtained from a unit in depositional contact on an immediately underlying volcanic unit have been assumed, based on stratigraphic proximity and petrography, to be derived solely from the volcanic unit thereby providing a maximum age constraint. Thus the dating of any sedimentary unit will rely upon a mixture of constraints made with and without inference, and these should be stated explicitly and supported by field observations and geologically reasonable arguments.

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Fig. 3. A schematic illustration of age relationships based on detrital, intrusive and extrusive ages and their relevance to their host unit and to encasing units. Unit B represents Lomagundi– Jatuli Event-bearing rocks.

We have used this framework to categorise the summary database of radio-isotopic dates reported in Table 1. 5. Chronological constraints on the Lomagundi–Jatuli Event Sections recording the LJE from each continent have been reviewed for location, formation/unit name, stratigraphic context, δ13C values and age data, including the type of mineral/rock dated and its relationship to strata with δ13C ≥ 5‰. Our emphasis has been on U–Pb minerals and Re–Os organic-rich sediments and the formations/units constrained with these robust chronometers are summarised in Table 1; a full listing of all radio-isotope dates associated with the LJE are included in Supplementary Table A1. References for radiometric ages and δ13C values are listed in Tables 1 and 2. The present day locations with longitude and latitude to all LJE-bearing sections discussed in this section are given in Supplementary Fig. A1; additional locality diagrams are included in the text where appropriate. Fig. 2 provides a plate reconstruction for c. 2100 Ma showing the location of the carbonate successions discussed in this section.

5.1. North America The LJE in North America is recorded in the Huronian and Marquette Range supergroups in the Great Lakes area, in Wyoming, USA, and in the Labrador Trough, Canada (Fig. 4A). In the Huronian Supergroup of the Great Lakes area, the Copper Cliff Rhyolite sits near the base and has been dated at 2452.5 ± 6.2 Ma (Ketchum et al., 2013). Overlying this, separated by several formations, LJE-bearing rocks crop out in the Gordon Lake Formation with δ13C values ≤ +8.2‰ (Bekker et al., 2006). The c. 2215 Ma (e.g., Noble and Lightfoot, 1992) Nipissing Intrusions cross cut the Gordon Lake Formation (Fig. 4B) providing a minimum age for the formation. In the Marquette Range, the Archaean basement is cross-cut by the Tilden Granite (2345 ± 20 Ma; Fig. 4B; Hammond, 1976) that is in-turn truncated by an unconformity and thus provides a maximum age for the overlying strata. The youngest detrital zircon from the overlying Enchantment Lake Formation, which lies directly above the unconformity, is 2317 ± 6 Ma (Vallini et al., 2006) in age, which is a maximum age constraint for this level and the overlying Kona Dolomite (separated from the Enchantment Lake

Table 1 Published ages relevant to the Lomagundi–Jatuli Event. Region Sierra Madre Mtns. Medicine Bow Mtns. San José GB Ashburton Basin Ashburton Basin Earaheedy Basin Yerrida Basin Langöy/Hinnöy Menominee Labrador Trough Labrador Trough Upper Group Earaheedy Basin Loch Maree Karelia Craton Kola Craton Karelia Craton Central Hurungwe Granite Central Earaheedy Basin Ashburton Basin Gweta Borehole Kola Craton Franceville Basin Transvaal Kola Craton Kola Craton Kola Craton Earaheedy Basin São Francisco Craton Franceville Basin San José GB – – Franceville Basin Sierra Madre Mtns. – Menominee Jatuli (tectofacies) São Francisco Craton Gweta Borehole Marquette Range – Labrador Trough San José GB Labrador Trough – Gogebic Range Hamersley Basin

Formation/Stratigraphy

Age (Ma)

(±)

Rock type

Method (a)

(b)

Reference to age

(c)

(d)

(e)

(f)

Fletcher Park Rhyolite Keystone Quartz Diorite Uruguayan Dyke swarm June Hill Volcanics June Hill Volcanics Mulgarra Sandstone Maralou Fm. Lødingen granite Hemlock Nimish Fm. Sokoman Fm. Jörn Granites Frere Fm. Ard Gneiss Flowerdale Schist Suisari Fm. Pilgujärvi Volcanic Fm. Medvezhgora Fm. Bremen Creek Granite Gneiss Hurungwe Granite Mills Lac Granite Yelma Fm. Wooly Dolomite Granitoid Kuetsjärvi Volcanic Fm. Francevillian A Väystäjä Fm. Bushveld Complex Il'mozero Sedimentary Fm. Kolosjoki Sedimentary Fm. Kolosjoki Sedimentary Fm. Chiall Fm. Itacolomi Group Francevillian D Isla Mala Granitic Suite Basement Amphibole Schist Francevillian D Meta-gabbro Hirsimaa Sturgeon Quartzite Koljola Fm. Sabará Group Paragneiss Enchantment Lake Fm. Petäjäskoski Mistamisk Fm. Paso Severino (rhyolite–dacite) Dunphy Fm. Siltstone Sunday Quartzite Sill in Meteorite Bore Member

1780 1781 1790 1795 1799 1816 1843 1873 1874 1878 1880 1891 1891 1907 1912 1969 1970 1976 1982 1997.5 2009 2017 2031 2039.2 2049 2050 2050 2054 2055.5 2056.6 2058 2058 2059 2072 2074 2076 2078 2083 2092 2106 2115 2116 2125 2125 2133 2140 2142 2146 2169 2206 2207 2208

6 7 5 7 8 26 10 2 9 2 2 7 8 3 7 18 5 9 5 2.6 7 15 6 1.4 28 30 8 2 2.3 0.8 2 14 58 29 6 5 8 6 9 8 5 24 4 6 11 11 4 7 4 9 5 10

Rhyolite Quartz diorite Gabbro dyke swarm Rhyolitic volcaniclastic Rhyolitic volcaniclastic Sandstone Dolerite sill Granite Rhyolite Syenite cobbles Carbonatite Granitoid Tuff Granodiorite Sediment Gabbro Felsic tuff Basalt Granite Granite Granite Sandstone Tuffaceous siltstone Granitoid Volcaniclastic Uraninite ore Felsic volcanic rocks Orthopyroxenite Meta-sediment Tuff Sandstone Turbidite Quartzite Sediment Granite Diabase intrusion Albite diabase intrusion Welded tuff Pegmatitic meta-gabbro Mafic pyroclastic rock Quartzite Metadiabase intrusion Sabará Group Paragneiss Sediment Gabbro Rhyolite Dacite Cramolet Lake gabbro sill Albite diabase intrusion Quartzite Mafic Sill

ID-TIMS ID-TIMS ID-TIMS SIMS SIMS SIMS SIMS ID-TIMS ID-TIMS ID-TIMS ID-TIMS ID-TIMS SIMS ID-TIMS SIMS TIMS (Re-Os) ID-TIMS SIMS ID-TIMS SIMS ID-TIMS SIMS SIMS ID-TIMS SIMS ID-TIMS ID-TIMS SIMS ID-TIMS ID-TIMS ID-TIMS SIMS ID-TIMS LA SIMS ID-TIMS ID-TIMS SIMS ID-TIMS SIMS SIMS ID-TIMS ID-TIMS ID-TIMS SIMS ID-TIMS ID-TIMS SIMS ID-TIMS ID-TIMS SIMS SIMS

Zrn Zrn Bdy Zrn Zrn Zrn Mnz Zrn, Ttn Zrn Zrn (F) Zrn Zrn (F) Zrn Zrn Zrn WR + Usp + Ilm Zrn Zrn Zrn (F) Zrn Zrn Zrn Zrn Zrn Zrn Urn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Bdy + Zrn Zrn + Ttn (F) Zrn Zrn Zrn Xtm Zrn (F) Zrn Zrn Xtm Zrn ? Zrn Zrn (F) Zrn + Ttn + Bdy(F) Xtm Bdy

Premo and Van Schmus (1989) Premo and Van Schmus (1989) Halls et al. (2001) Wilson et al. (2010) Evans et al. (2003) Halilovic et al. (2004) Rasmussen and Fletcher (2002) Corfu (2004) Schneider et al. (2002) Findlay et al. (1995) Chevé and Machado (1988) Wilson et al. (1987) Rasmussen et al. (2012) Park et al. (2001) Whitehouse et al. (1997) Puchtel et al. (1999) Hanski et al. (1990) Puchtel et al. (1998) Goldich and Fischer (1986) McCourt et al. (2001) Holm et al. (2005) Halilovic et al. (2004) Müeller et al. (2005) Majaule et al. (2001) Melezhik et al. (2007) Gancarz (1978) Perttunen and Vaasjoki (2001) Scoates and Friedman (2008) Martin et al. (2013) Martin et al. (2013) Melezhik et al. (2007) Halilovic et al. (2004) Machado et al. (1996) Bouton et al. (2009a, 2009b) Hartmann et al. (2000) Buchan et al. (1996) Silvennoinen (1991) Horie et al. (2005) Premo and Van Schmus (1989) Karhu et al. (2007) Vallini et al. (2006) Pekkarinen and Lukkarinen (1991) Machado et al. (1992) Mapeo et al. (2001) Vallini et al. (2006) Kyläkoski et al. (2013) Clark (1984) Santos et al. (2003) Rohon et al. (1993) Silvennoinen (1991) Vallini et al. (2006) Müeller et al. (2005)

E I I E E D D I E D I I E I D I E I I I I D E I D

min min min min min no no min min no min min min min no min min min min no min no min no min max min min min min min no no no no no min min min Syn no Max No No no min syn no min min no no

P P P P P P P P P P P P P P U P P P P U P P P U P U P P U P P P U U U U U U U S U U U U U S U U U S U U

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52

E I D E D D D D I I I E I E M E D D M I E E I I M I

(continued on next page)

247

Wyoming Wyoming Uruguay Pilbara Craton Pilbara Craton Yilgarn Craton Yilgarn Craton Lofoten Great Lakes Canada Canada Kalix GB Yilgarn Craton Lewisian Lewisian Onega Basin Pechenga GB Onega Basin Great Lakes Zimbabwe Great Lakes Yilgarn Craton Pilbara Craton Zimbabwe Pechenga GB Gabon Peräpohja SB South Africa I–V GB Pechenga GB Pechenga GB Yilgarn Craton Brazil Gabon Uruguay Great Lakes Kuusamo SB Gabon Wyoming Peräpohja SB Great Lakes Kiihtelysvaara Brazil Zimbabwe Great Lakes Peräpohja SB Canada Uruguay Canada Kuusamo SB Great Lakes Pilbara Craton

A.P. Martin et al. / Earth-Science Reviews 127 (2013) 242–261

North America North America South America Australia Australia Australia Australia Norway North America North America North America Sweden Australia Scotland Scotland Russia Russia Russia North America Africa North America Australia Australia Africa Russia Africa Finland Africa Russia Russia Russia Australia South America Africa South America North America Finland Africa North America Finland North America Finland South America Africa North America Finland North America South America North America Finland North America Australia

248

Table 1 (continued) Region Pilbara Craton Peräpohja SB Great Lakes Peräpohja SB Zimbabwe Great Lakes South Africa Great Lakes South Africa Great Lakes Kuusamo SB Lewisian South Africa Wyoming Peräpohja SB I–V GB South Africa Rajasthan I–V GB I–V GB Onega Basin Pilbara Craton Pilbara Craton Great Lakes Wyoming Pilbara Craton South Africa Lewisian South Africa Pechenga GB Pechenga GB Pechenga GB South Africa South Africa Great Lakes Pilbara Craton Zimbabwe South Africa Pilbara Craton Brazil Great Lakes Canada Wyoming Canada Uruguay

Ashburton Basin – Huronian – Chimbadzi Hill Menominee Transvaal Marquette Range Transvaal – – – Transvaal Medicine Bow Mtns. – Kola Craton Griqualand West Mewar Gneissic Kola Craton Kola Craton Karelia Craton Hamersley Basin Hamersley Basin Thessalon Fm. Sierra Madre Mtns. Hamersley Basin Griqualand West – Transvaal Kola Craton Kola Craton Kola Craton Griqualand West Griqualand West – Hamersley Basin Basement Transvaal Hamersley Basin São Francisco Craton Gogebic Range Labrador Trough Sierra Madre Mtns. Labrador Trough San José GB

Formation/Stratigraphy

Age (Ma)

(±)

Rock type

Method (a)

(b)

Reference to age

(c)

(d)

(e)

(f)

Cheela Springs Basalt Palokivalo Fm. Nipissing Intrusions Laurila Sill Chimbadzi Hill layered complex Sturgeon Quartzite Rooihoogte Fm. Enchantment Lake Fm. Timeball Hill Fm. Tilden Granite Conglomerate Scourie Dyke Swarm Duitschland Fm. Baggot Rocks Granite Elijärvi granite Polisarka Sedimentary Fm. Makganyene Delwara Fm. Imandra lopolith Seidorechka Volcanic Fm. Burakovka Pluton Woongarra Rhyolite Weeli Wolli Fm. Copper Cliff Rhyolite Magnolia Fm. Brockman Iron Fm. Kuruman Fm. Granitic Pegmatite Penge Pana Tundra Intrusion Monche Pluton General'skaya Intrusion Gamohan Gamohan McGrath Gneiss Wittenoom Fm. Great Dyke Oak Tree Fm. Marra Mamba Iron Fm. Moeda Sunday Quartzite Basement Basement Basement Basement

2209 2215 2217 2221 2262 2306 2316 2317 2324 2345 2405 2418 2424 2429 2433 2434 2436.2 2440 2441 2442 2449 2449 2449 2452.5 2451 2454 2460 2480 2480 2501.5 2504.4 2505 2516 2521 2556 2561 2579 2582 2597 2606 2647 2654 2683 2692 2762

15 0 4 5 2 9 7 6 17 20 6 7 12 4 4 1.2 6.6 8 1.6 1.7 1.1 3 3 6.2 9 3 5 1 6 1.7 1.5 1.6 4 3 10 8 3 15 5 47 5 5 6 9 8

Volcaniclastic breccia Albite sills Diabase Albite diabase Troctolite Quartzite Carbonaceous shale Sediment Quartzite Granite Quartz porphyry Bronzite picrite Sediment (diamictite) Granite Granite Tuff Sediments Granitoid Granophyre Felsic subvolcanic unit Gabbro-norite Rhyolite lava flow Rhyolite tuff Rhyolite Quartz meta-conglomerate Tuffaceous mudrocks Felsic Tuff Granite – Mafic layered intrusion Mafic layered intrusion Mafic layered intrusion Meta-sediment – Basement Tuff Websterite Felsic tuff Tuff Quartzite Quartzite Granite pegmatite Orthogneiss Foliated tonalite Meta-conglomerate

SIMS ID-TIMS ID-TIMS ID-TIMS ID-TIMS SIMS TIMS (Re-Os) SIMS SIMS ? ID-TIMS U–Pb SIMS ID-TIMS ID-TIMS ID-TIMS SIMS SIMS ID-TIMS ID-TIMS ID-TIMS U–Pb U–Pb ID-TIMS ID-TIMS SIMS SIMS ID-TIMS SIMS ID-TIMS ID-TIMS ID-TIMS SIMS ID-TIMS ID-TIMS SIMS SIMS SIMS SIMS LA SIMS ID-TIMS ID-TIMS ID-TIMS SIMS

Zrn Zrn, Ttn, Bdy Bdy Bdy Bdy Zrn Py Zrn Zrn ? Zrn (F) Bdy Zrn Zrn Zrn Zrn Zrn Zrn Zrn Bdy Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn Zrn (F) Zrn Zrn Zrn Zrn Zrn Zrn Zrn Mnz + Zrn Zrn Zrn (F) Zrn

Martin et al. (1998) Perttunen (1991) Noble and Lightfoot (1992) Perttunen and Vaasjoki (2001) Manyeruke et al. (2004) Vallini et al. (2006) Hannah et al. (2004) Vallini et al. (2006) Dorland (2004) Hammond (1976) Silvennoinen (1991) Heaman and Tarney (1989) Dorland (2004) Premo and Van Schmus (1989) Perttunen and Vaasjoki (2001) Brasier et al. (2013) Moore et al. (2012) Wiedenbeck et al. (1996) Amelin et al. (1995) Amelin et al. (1995) Amelin et al. (1995) Barley et al. (1997) Barley et al. (1997) Ketchum et al. (2013) Premo and Van Schmus (1989) Pickard (2002) Pickard (2003) Corfu et al. (1994) Nelson et al. (1999) Amelin et al. (1995) Amelin et al. (1995) Amelin et al. (1995) Altermann and Nelson (1998) Sumner and Bowring (1996) Holm et al. (2005) Trendall et al. (1998) Armstrong and Wilson (2000) Martin et al. (1998) Trendall et al. (1998) Machado et al. (1996) Vallini et al. (2006) Mortensen and Percival (1987) Premo and Van Schmus (1989) Mortensen and Thorpe (1987) Hartmann et al. (2001)

E I I I I D A D D I D I D I I E D I I E I E E E D E E I D I I I D I I E I E E D D I I I D

no min min min no max max max max max max max max max max max max max no max max max max max max max max max max max max max max max max max max max max max max max max max max

U U U S U Pr Pr U Pr U U Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr Pr

53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 93 94 95 96 97

(a)All methods have utilised the U–Pb decay schemes and reported ages are concordant (or upper intercept) 207Pb/206Pb dates, except data 15 and 58 that have utilised the Re–Os decay scheme. (b)Sample type. (c)Nature of the dated unit. D = detrital. I = intrusive. E = extrusive. A = authogenic/diagenetic. M = metamorphic overgrowth. (d)Nature of the age constraint in relation to the local Lomagundi–Jatuli Event-bearing unit(s) min = minimum age constraint. max = maximum age constraint. syn = contemporaneous age constraint. no = no age constraint. (e)Stratigraphic constraint in relation to the Lomagundi–Jatuli Event. P = post. S = syn. Pr = pre. U = Not assigned a category. (f)Number corresponds to dates used in the figures and text. I–V (Imandra–Varzuga); GB (Greenstone Belt); SB (Schist Belt); Mtns. (Mountains); Fm. (Formation); ID-TIMS (Isotope Dilution Thermal Ionisation Mass Spectrometry); SIMS (Secondary Ion Mass Spectrometer); LA (Laser Ablation Inductively Coupled Plasma Mass Spectrometry); Zrn (zircon); WR (whole rock); Bdy (baddeleyite); Usp (ulvöspinel) Ilm (ilmenite); Rt (rutile); Py (pyrite); Mnz (monazite); Urn (uraninite); Xtm (xenotime); (F) = Fraction. All mineral abbreviations follow Whitney and Evans (2010).

A.P. Martin et al. / Earth-Science Reviews 127 (2013) 242–261

Australia Finland North America Finland Africa North America Africa North America Africa North America Finland Scotland Africa North America Finland Russia Africa India Russia Russia Russia Australia Australia North America North America Australia Africa Scotland Africa Russia Russia Russia Africa Africa North America Australia Africa Africa Australia South America North America North America North America North America South America

A.P. Martin et al. / Earth-Science Reviews 127 (2013) 242–261 Table 2 References to δ13C values used in Figs. 4–11. a

Reference

Region

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

Bekker et al. (2006) Melezhik et al. (1997) Bekker et al. (2003a) Melezhik et al. (1999) Melezhik et al. (2007) Melezhik and Fallick (1996) Karhu (2005) Karhu (1993) Préat et al. (2011) Frauenstein et al. (2009) Bau et al. (1999) Melezhik and Fallick (2010) Bekker et al. (2001) Master et al. (2010) Lindsay and Brasier (2002) Bekker et al. (2003b) Maheshwari et al. (2010) Kump et al. (2011) Planavsky et al. (2012) Krupenik et al. (2011)

Great Lakes area Labrador Trough Wyoming Onega Basin Pechenga Imandra–Varzuga Peräphoja Kuusamo Gabon Griqualand West Griqualand West Transvaal Transvaal Zimbabwe Australia Brazil Uruguay Onega Basin Zimbabwe Onega Basin

249

Chevé and Machado, 1988) and Nimish (1878 ± 2 Ma; Findlay et al., 1995) formations (Fig. 5A). In Wyoming (Fig. 4A), LJE-bearing sedimentary rocks are recorded in the Nash Fork Formation (δ13C values ≤ +28‰; Bekker et al., 2003a) of the Libby Creek Group in the Medicine Bow Mountains (Fig. 5B). The Group rests on the 2429 ± 4 Ma (Premo and van Schmus, 1989) Baggot Rocks granite and is cross-cut by the 1781 ± 7 Ma (Premo and van Schmus, 1989) Keystone quartz diorite intrusions (Fig. 5B); combined, these provide maximum and minimum age constraints to the Nash Fork Formation. In the nearby Sierra Madre Mountains, the Snowy Pass Group (Fig. 5B) rests on a 2683 ± 6 Ma basement (Premo and van Schmus, 1989). Detrital zircons dated at 2451 ± 9 Ma (Premo and van Schmus, 1989) in the Magnolia Formation at the base of that group provide a maximum age constraint for the overlying Slaughterhouse Formation (Fig. 5B), which contains the LJE (δ13C values ≤ +16.6‰; Bekker et al., 2003a). The Slaughterhouse Formation is overlain structurally by the 1780 ± 6 Ma Fletcher Park rhyolite porphyry and is intruded by a 2092 ± 9 Ma pegmatitic meta-gabbro (Premo and van Schmus, 1989).

(a)Keyed to numbers in Figs. 4–11.

5.2. Russia

Formation by the intervening Mesnard Quartzite) that contains the LJE (δ13C values ≤ +9.5‰; Bekker et al., 2006; Fig. 4B). Xenotime overgrowths have been dated at 2133 ± 11 Ma (Vallini et al., 2006) in the Enchantment Lake Formation but provide no age constraint to the overlying Kona Dolomite. In the Menominee and Iron RiverCrystal Falls Ranges, the Randeville Dolomite has normal δ13C values (≤+3.1‰; Bekker et al., 2006). It overlies the Sturgeon Quartzite in which the youngest detrital zircon age is given at 2306 ± 9 Ma (Vallini et al., 2006). Xenotime overgrowths on zircon in the Randeville Dolomite have been dated at 2115 ± 5 Ma (Vallini et al., 2006) giving a minimum age to this formation, which is overlain by the Hemlock Formation dated at 1874 ± 9 Ma (Schneider et al., 2002; Fig. 4B). In the Gogebic Range, detrital zircons from the Sunday Quartzite yield a maximum age at 2647 ± 9 Ma (Vallini et al., 2006; Fig. 4B). The McGrath Gneiss from the basement of the Marquette Range Supergroup in central Minnesota has a crystallisation age at 2556 ± 10 Ma (Holm et al., 2005; Fig. 4B). This basement is cross-cut by an igneous body dated at 2076 ± 5 Ma (Buchan et al., 1996) that is in turn truncated by an unconformity (Fig. 4B) providing a maximum age to the overlying strata. The Mille Lacs Group records δ13C values ≤ +2.5‰ (Bekker et al., 2006) and is cross-cut by both the Mills Lac Granite (2009 ± 7 Ma; Holm et al., 2005) and Bremen Creek Granite Gneiss (1982 ± 5 Ma; Goldich and Fischer, 1986). In the Labrador Trough (Fig. 4A), Palaeoproterozoic rocks rest unconformably on Archaean basement dated at 2692 ± 9 Ma (Mortensen and Thorpe, 1987) and 2654 ± 5 Ma (Mortensen and Percival, 1987; Fig. 5A). In the Eastern zone of the Labrador Trough, the LJE is recorded in the Dunphy Formation (δ13C = + 16‰; Melezhik et al., 1997). This is overlain by Mistamisk Formation rhyolites, dated at 2142 ± 4 Ma (Clark, 1984), then the Denault and Abner formations with δ13C ≤ + 4‰ (Melezhik et al., 1997; Fig. 5A). In the Western zone, the LJE is recorded in the Portage Formation (δ13C =+ 6.1‰; Melezhik et al., 1997) that is intruded by the gabbroic Cramolet Lake Sill (Fig. 5A) dated at 2169 ± 4 Ma (Rohon et al., 1993). This is a minimum age constraint for the Portage Formation, but cannot constrain the age of the overlying LJE-bearing Alder (+ 8 ≤ δ13C ‰ ≤ + 12) and Uvé (+ 5 ≤ δ13C ‰ ≤ + 8) formations (Melezhik et al., 1997). Above these units, δ13C values are normal in the Denault-Abner (+2 ≤ δ13C ‰ ≤ +4; Melezhik et al., 1997) and Doly (δ13C ‰ = −0.6; Melezhik et al., 1997) formations, and post-LJE rocks crop out in cycle 2 in the Sokoman (1880 ± 2 Ma;

On the Kola Peninsula, LJE-bearing rocks are known from the North Transfennoscandian Greenstone Belt that stretches discontinuously for over c. 1000 km across northwest Russia and adjacent parts of northern Norway and Finland (Fig. 6A). In Russia, the best-studied sections are the Pechenga and the Imandra–Varzuga Greenstone Belts, and the Onega Basin (Fig. 6A). The Burakovka Pluton (Fig. 6B), dated at 2449 ± 1.1 Ma (Amelin et al., 1995), cross-cuts the Archaean basement and provides a maximum age for the unconformably overlying Onega Basin rocks. The LJE is recorded in the Tulomozero Formation (Fig. 6B), with δ13C values between +3.5‰ and +17.2‰ (Melezhik et al., 1999; Brasier et al., 2011). Above this is the Zaonega Formation, with δ13C between −20 and +10‰ (Krupenik et al., 2011; Kump et al., 2011; Fig. 6B). A 207Pb/206Pb date of 1976 ± 9 Ma has been found on an igneous rock in the Onega Basin succession, however its stratigraphic position is uncertain and has been assigned to either the Medvezhegora Formation or the Suisari Formation (Puchtel et al., 1998), although a Re–Os isochron age of 1969 ± 18 Ma (Puchtel et al., 1999) on a Suisari Formation suggests a correlation with the latter (Fig. 6B). The Pechenga Greenstone Belt rests on Archaean basement that is cross-cut by multiple intrusions, including the General'skaya Intrusion dated at 2505 ± 1.6 Ma (Amelin et al., 1995; Fig. 6C); this places a maximum age on the Pechenga Greenstone Belt. The Kuetsjärvi Sedimentary Formation records the LJE with δ13C values of +9‰ that decline up section to +6‰ (Melezhik et al., 2007). Overlying this is the Kolosjoki Volcanic and Sedimentary formations. The latter has δ13C values between +1‰ and +2.5‰ (Melezhik et al., 2007; Fig. 6C), indicating that the LJE had terminated prior to its deposition. A volcanoclastic pebble conglomerate within the Kuetsjärvi Volcanic Formation has yielded two zircon grains with an age at 2049 ± 29 Ma (Fig. 6C) whereas detrital zircons obtained from the basal units of the Kolosjoki Sedimentary Formation (which consist of monomict, volcaniclastic, conglomerate sourced directly from the underlying Kuetsjärvi Volcanic Formation) have yielded ages of 2058 ± 2 Ma (Melezhik et al., 2007). A tuff, also from near the base of the Kolosjoki Sedimentary Formation, is dated at 2056.6 ± 0.8 Ma (Martin et al., 2013; Fig. 6C) and is sandwiched between units having normal δ13C values. Further, the Pilgujärvi Volcanic Formation (Fig. 6C), which defines the top of the North Pechenga Group, contains felsic tuffs, one of which has yielded a zircon dated at 1970 ± 5 Ma (Hanski et al., 1990) and providing a maximum age for this formation. The Imandra–Varzuga Greenstone Belt lies unconformably on Archaean rocks (Fig. 6D). There, the base of the Strelna Group rests on the 2504.4 ± 1.5 Ma Monche Pluton and the 2501.5 ± 1.7 Ma Pan Tundra Intrusion (Amelin et al., 1995). An intrusion cross cuts these

250

A.P. Martin et al. / Earth-Science Reviews 127 (2013) 242–261

Fig. 4. (A) Map of key North American Lomagundi–Jatuli Event localities. (B) Stratigraphic sections in the Great Lakes area North America (after Bekker et al., 2006). Age references are given in Table 1. References to δ13C are given in Table 2. On = Ontario. Mi = Michigan. Wi = Wisconsin. Mn = Minnesota.

rocks that may be a feeder to Seidorechka Volcanic Formation lava flows and is dated at 2442 ± 1.7 Ma and is spatially and temporally associated with the Imandra lopolith dated at 2441 ± 1.6 Ma (Amelin et al., 1995; Fig. 6D). Zircons from a tuff in the overlying Polisarka Sedimentary Formation have been dated at 2434 ± 1.2 Ma (Brasier et al., 2013); the eruption of the tuff is interpreted as contemporaneous with deposition of the encasing sediments. The overlying Umba Sedimentary Formation contains the LJE as indicated by carbonate rocks with δ13C ≤ +6.67‰ (Melezhik and Fallick, 1996; Fig. 6D). Above this are the Umba Volcanic and Il'mozero Sedimentary formations. The latter has carbonate rocks with δ13C values between −0.83‰ and +2.36‰ (Melezhik and Fallick, 1996) and contains detrital zircon that yield a minimum date of deposition at 2055.5 ± 2.3 Ma (Martin et al., 2013). 5.3. Finland Finland has several Palaeoproterozoic belts where LJE-bearing rocks crop out and these are reviewed in detail in Karhu (1993). Here, the three areas with the best chronological constraints are highlighted: the Peräphoja Belt, the Kuusamo Schist Belt and the Kiihtelysvaara area (Fig. 6A). The basement of the Peräphoja Belt (Fig. 7A) is cross-

cut by the Elijärvi Granite dated at 2433 ± 4 Ma (Perttunen and Vaasjoki, 2001) and is a maximum age constraint to the overlying supracrustal rocks. The Sompujärvi, Rankaus and base of the Palokivalo formations are cross-cut by the Laurila Sill dated at 2221 ± 5 Ma (Perttunen and Vaasjoki, 2001). The top of the Palokivalo Formation contains carbonate rocks with δ13C values between +5‰ and +11‰ (Karhu, 2005), and is cross-cut by diabase sills dated at c. 2215 Ma (Perttunen, 1991). These sills and the Laurila Sill provide a minimum age to the Palokivalo Formation. Recent work has recognised and defined a new unit, the Petäjäskoski Formation (at the base of the Jouttiaapa Formation) that is cross-cut by a gabbro-sill yielding zircon dated at 2140 ± 11 Ma (Kyläkoski et al., 2013). This age is a minimum age to the Petäjäskoski Formation. Further, the laterally restricted Hirsimaa Formation (between the more extensive Tikanmaa and Rantamaa formations) contains mafic pyroclastic rocks from which zircon have yielded a preliminary date of 2106 ± 8 Ma (Karhu et al., 2007). The Kvartsimaa and Rantamaa formations record the LJE (Fig. 7A) with the former having δ13C values ≤ +10‰ and the latter showing δ13C values of +10‰ that decline up section to +2‰ (Karhu, 2005). The overlying Väystäjä Formation has normal δ13C values of −1‰ to 0‰ (Karhu, 2005) and contains syn-sedimentation felsic volcanic rocks dated at 2050 ± 8 Ma (Perttunen and Vaasjoki, 2001; Fig. 7A).

A.P. Martin et al. / Earth-Science Reviews 127 (2013) 242–261

251

Fig. 5. Key Lomagundi–Jatuli Event localities from North America. See Fig. 4 for the locations and legend. Age references are given in Table 1. References to δ13C are given in Table 2. (A) Stratigraphic section from the Labrador Trough, Canada, after Rohon et al. (1993) and Melezhik et al. (1997). (B) Stratigraphic section from Wyoming, America, after Bekker et al. (2003a).

In the Kuusamo Belt (Fig. 7B), the basal unit, the Conglomerate Formation, has yielded detrital zircons that provide a maximum age constraint of 2405 ± 6 Ma (Silvennoinen, 1991). The LJE is recorded in the overlying Sericite Schist Formation, with δ13C ≤ +8‰ and the Siltstone Formation, with δ13C of +12‰ (Karhu, 1993). These units are cross-cut by a diabase intrusion dated at 2206 ± 9 Ma (Silvennoinen, 1991), which provides a minimum age for the two formations. A second generation of diabase intrusions, dated at 2078 ±8 Ma (Silvennoinen, 1991), cross-cuts the Rukatuntari Quartzite and may be the feeder to the Amphibole Schist Formation above (Karhu, 1993). The primary nature of the dated material in these intrusions has been questioned (Hanski et al., 2001), though the dyke has yet to be re-analysed. The Dolomite and Limestone–Dolomite formations at the top of the Kuusamo Schist Belt (Fig. 7B) also record the LJE, having δ13C values of +11.5‰ and +7‰, respectively (Karhu, 1993). In the Kiihtelysvaara area, the LJE-bearing rocks of the Hypiä Group are the Viistola Formation, with δ13C between +4‰ and +12‰, and the base of the Petäikkö Formation, having δ13C ≤ +6‰ that decreases up section to +0.5‰ (Karhu, 1993; Fig. 7C). Underlying the Hypiä Group is a meta-diabase intrusion feeding lavas of the Koljola Formation dated at 2116 ± 24 Ma (Pekkarinen and Lukkarinen, 1991; Fig. 7C). 5.4. Africa LJE-bearing units occur in South Africa, Zimbabwe and Gabon (Fig. 8). In South Africa, LJE-bearing rocks occur in the Transvaal and

Griqualand West areas (Fig 8A). In the former, felsic tuffs in the Oak Tree and Penge formations have been dated at 2582 ± 15 Ma (Martin et al., 1998) and 2480 ± 6 Ma (Nelson et al., 1999), respectively (Fig. 8B). Overlying the Penge Formation is the Duitschland Formation with δ13C values of ≤ +10‰ (Bekker et al., 2001). However, Frauenstein et al. (2009) have questioned the primary nature of the δ13C values. They have shown that δ13C values are normal in carbonate rocks having elevated Sr values (~900 ppm) and only where Sr values are anomalously low is δ13C ≥ +4‰, suggesting to them that the C isotopes have undergone post-depositional alteration. The youngest, concordant detrital zircon age from a diamictite in the Duitschland Formation has been determined at 2424 ± 12 Ma (Dorland, 2004): this is a maximum age for the level sampled and for overlying strata. The Rooihoogte Formation is considered to be correlative to the Duitschland Formation (Hannah et al., 2004) and the former contains diagenetic pyrite dated by the Re–Os technique at 2316 ± 7 Ma (Hannah et al., 2004; Fig. 8B). This is interpreted as close to the depositional age of the Rooihoogte Formation and, hence, that of the Duitschland Formation. This date is one of the few to robustly constrain the depositional age of LJE-bearing rocks. However, if the arguments of Frauenstein et al. (2009) are accepted, then the 2316 ± 7 Ma age must pre-date the LJE and becomes one of the youngest dates for pre LJE-bearing units. The Re–Os date is supported by unpublished 207Pb/ 206 Pb data from the conformably overlying Timeball Hill Formation where detrital zircons yielded a maximum age of 2324 ± 17 Ma (as noted in Dorland, 2004). Overlying the Duitschland Formation, and separated by several units, is the Silverton Formation with δ13C

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Fig. 6. Key Lomagundi–Jatuli Event localities from Russia. See Fig. 4 for the legend. Age references are given in Table 1. References to δ13C are given in Table 2. (A) Map of key Russian and Finnish Lomagundi–Jatuli Event localities. (B) Stratigraphic section from the Onega Basin after Melezhik et al. (1999). (C) Stratigraphic section from the Pechenga Greenstone Belt after Melezhik and Sturt (1994). (D) Stratigraphic section from the Imandra–Varzuga Greenstone Belt after Melezhik and Sturt (1994).

values ≤ +10‰ (Frauenstein et al., 2009; Fig. 8B). Stratigraphically above this, δ13C values return to normal (−3.3 ≤ δ13C ‰ ≤ −0.5; Melezhik and Fallick, 2010) in the Houtenbeck Formation (Fig. 8B). Above the Houtenbeck Formation, zircons in the Merensky Reef of the Bushveld Complex are dated at 2054 ± 2 Ma (Scoates and Friedman, 2008). This provides a minimum age to the Silverton Formation (Fig. 8B). In the Griqualand West area, the Lucknow Formation records the LJE with δ13C ≤ +10‰ (Frauenstein et al., 2009) and normal δ13C values are recorded in the underlying Mooidrai (+0.5 ≤ δ13C ‰ ≤ +2) and Rooinekke (−8 ≤ δ13C ‰ ≤ −1) formations (Bau et al., 1999). At the

base of the Griqualand West stratigraphy an igneous zircon date in the Gamohaan Formation yields an age of 2521 ± 3 Ma (Sumner and Bowring, 1996), supported by the youngest detrital zircon age from the same formation dated at 2516 ± 4 Ma (Altermann and Nelson, 1998). Overlying this, a felsic tuff in the Kuruman Formation has been dated at 2460 ± 5 Ma (Pickard, 2003), and in the overlying Makganyene Formation detrital zircon has yielded a 2436.2 ± 6.6 Ma age (Moore et al., 2012; Fig. 8B). In Gabon, the FB (δ13C ≤ +6.47‰) and FC (δ13C ≤ +9.74‰) formations in the Franceville Basin (Préat et al., 2011; Fig. 8C) record

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Fig. 7. Key Lomagundi–Jatuli Event localities from Finland. See Fig. 4 for the legend and Fig. 6A for localities. Age references are given in Table 1. References to δ13C are given in Table 2. Stratigraphic sections from the (A) Peräphoja Belt, (B) Kuusamo Schist Belt and (C) Kiihtelysvaara area after Karhu (1993).

the LJE. The older FA Formation contains uraninite ore dated at 2050 ± 30 Ma (Gancarz, 1978; Fig. 8C) and the younger FD Formation, which is unconformable on the FC, contains detrital zircon dated at 2072 ± 29 Ma, as well as zircon from a welded tuff dated at 2083 ±6 Ma (Bouton et al., 2009a, b; Fig. 8C). In Zimbabwe, δ13C values N + 5‰ are recorded in the Lomagundi and Deweras groups of the Magondi Mobile Belt (Fig. 9). Zimbabwe's Great Dyke, which cross cuts basement rocks and is dated at 2579 ± 3 Ma (Armstrong and Wilson, 2000), provides a maximum age constraint for these units (Fig. 9). At the base of the Magondi Mobile Belt, a troctolite of the Chimbadzi Hill Layered Complex is dated at 2262 ± 2 Ma (Manyeruke et al., 2004; Fig. 9) and is considered by some workers as the earliest Palaeoproterozoic magmatism on the Zimbabwe Craton and the intrusive equivalent of the base of the Deweras Group (Manyeruke et al., 2004). However, this linkage is ambiguous because nowhere is the contact observed between the troctolite and base of the Deweras Group (Master et al., 2010). In the Deweras Group, the LJE is recorded in the Mangula (+7 ≤ δ13C ‰ ≤ +14.6) and Norah (+13.1 ≤ δ13C ‰ ≤ +16.6) formations (Master et al., 2010; Fig. 9). The Lomagundi Group unconformably overlies the Deweras Group, with the Mcheka Formation recording δ13C values ≤ +11.9‰ (Planavsky et al., 2012; Fig. 9). Detrital zircons from paragneiss of the Copper Queen Formation provide an age of 2125 ± 6 Ma (Mapeo et al., 2001), though its stratigraphic position is described as ‘uncertain’ (Mapeo et al., 2001). A granitoid from Kuba Island, thought to be related to the Magondi Mobile Belt, is dated at 2039.2 ± 1.4 Ma (Majaule et al., 2001), but its stratigraphic position is uncertain. Finally, the Hurungwe Granite that cross-cuts the Pirwiri and Lomagundi groups is dated at 1997.5 ± 2.6 Ma (McCourt et al., 2001). 5.5. Australia The LJE signal is recorded in the Juderina Formation (δ13C values ≤ +11.9‰) of the Yerrida Basin, on the northern margin of the

Yilgarn Craton (Fig. 10A), Western Australia (Lindsay and Brasier, 2002; Fig. 10A, B). The Juderina Formation is overlain by the Johnson Cairn Formation recording normal δ13C values (−0.4‰; Lindsay and Brasier, 2002). At the top of the Yerrida Basin succession, a dolerite sill in the Maralou Formation is dated at 1843 ± 10 Ma (Rasmussen and Fletcher, 2002). Rocks in the Earaheedy, Hamersley and Ashburton basins occur adjacent to the Yerrida Basin (Fig. 10A). These three basins do not record the LJE but contain several dated units relevant to constraining the timing of the LJE in the Yerrida Basin. Detrital zircons provide maximum ages of 2017 ± 15 Ma (Halilovic et al., 2004) for the Yelma Formation (−1 ≤ δ13C ‰ ≤ +3; Lindsay and Brasier, 2002) at the base of the Earaheedy succession, 2058 ± 14 Ma (Halilovic et al., 2004) for the overlying Chiall Formation, and 1816 ± 26 Ma (Halilovic et al., 2004) for Mulgarra Sandstone at the top of the succession. Zircons from a tuff at the base of the Frere Formation yield an age of 1891 ± 8 Ma (Rasmussen et al., 2012) that is contemporaneous with deposition. In the Hamersley Basin (Fig. 10A), normal δ13C values (Lindsay and Brasier, 2002) are recorded in the Kazput Formation (− 6.5 ≤ δ13C‰ ≤ + 2), Meteorite Bore member (− 1.5 ≤ δ13C‰ ≤ + 1) and the Wittenoom Formation (−2 ≤ δ13C‰ ≤ +1.5). A cross-cutting mafic sill in the Meteorite Bore Member is dated at 2208 ± 10 Ma (Müeller et al., 2005) providing a minimum age to the latter two units. Given the conformable contacts between formations of the Turee Creek Group, the 2208 ± 10 Ma age may also be interpreted as a minimum age for the Kazput Formation. The other radiometric ages from the Hamersley Basin include 2597 ± 5 Ma (Trendall et al., 1998) from the Marra Mamba Formation, 2561 ± 8 Ma (Trendall et al., 1998) from the Wittenoom Formation, 2454 ± 3 Ma (Pickard, 2002) from the Brockman Iron Formation, 2449 ± 4 Ma (Barley et al., 1997) from the Weeli Wolli Formation and 2449 ± 3 Ma (Barley et al., 1997) from the Woongarra Rhyolite (Fig. 10A). Near the base of the Ashburton Basin, a volcaniclastic-rich breccia in the Cheela Springs Basalt is dated at 2209 ± 15 Ma (Martin et al., 1998; Fig. 10A) which has been attributed as either the age of the Cheela

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Fig. 8. Location and stratigraphic sections for Africa. See Fig. 4 for the legend. Age references are given in Table 1. References to δ13C are given in Table 2. (A) Map of key African Lomagundi– Jatuli Event localities. (B) Stratigraphic section from South Africa after Frauenstein et al. (2009). (C) Stratigraphic section from the Franceville Basin, Gabon, after Préat et al. (2011).

Springs Basalt (Martin et al., 1998) or a provenance age (Müeller et al., 2005). The overlying Wooly Dolomite records typically normal δ13C values (≤+2‰) with the exception of one datum with δ13C value of +6‰ (Lindsay and Brasier, 2002). A syn-sedimentary tuff close to the base of the Wooly Dolomite gives an age of 2031 ± 6 Ma (Müeller et al., 2005; Fig. 10A). At the top of the Ashburton Basin, two dates on rhyolite lava flows at 1799 ± 8 Ma (Evans et al., 2003) and 1795 ± 7 Ma (Wilson et al., 2010) provide a minimum age for the basin (Fig. 10A). 5.6. South America The LJE is recorded in the Minas Supergroup from the Quadilátero Ferrífero of the São Francisco Craton (Brazil) and from the San José Greenstone Belt on the Rio de la Plata Craton (Uruguay; Fig. 11A, B). Detrital zircons from the Moeda Formation at the base of the Minas succession yield a maximum age of 2606 ± 47 Ma (Machado et al.,

1996) for this and overlying units. In the overlying Piracicaba Group the Cercadinho (+3 ≤ δ13C ‰ ≤ +5.5) and Fecho do Funil formations (+5.5 ≤ δ13C ‰ ≤ +7.5) record the LJE (Bekker et al., 2003b). Stratigraphically above the Piracicaba Group is the Sabará Group dated at 2125 ± 4 Ma (Machado et al., 1992), which in turn is overlain by the Itacolomi Group where detrital zircons yield an age of 2059 ± 58 Ma (Machado et al., 1996). The San José Greenstone Belt (Fig. 11C) sits on Archaean basement for which the youngest, published age is 2762 ± 8 Ma (Hartmann et al., 2001). The Paso Severino Formation rests unconformably on the basement (Maheshwari et al., 2010) and contains carbonate rocks that record the LJE (+3.5 ≤ δ13C ‰ ≤ +11.6; Maheshwari et al., 2010). These are overlain by a dacite lava flow dated at 2146 ± 7 Ma (Santos et al., 2003). The Isla Mala granite cross-cuts the Paso Severino Formation and is dated at 2074 ± 6 Ma (Hartmann et al., 2000). The Uruguayan Tholeiitic Dyke swarm dated at 1790 ± 5 Ma (Halls et al., 2001) also cross-cut the Paso Severino Formation (Fig. 11C).

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Fig. 9. Stratigraphic section from the Magondi Mobile Belt, Zimbabwe, after Master et al. (2010). The Deweras Group shown is from the Northern facies area. See Fig. 4 for the legend and Fig. 8A for localities. Age references are given in Table 1. References to δ13C are given in Table 2.

5.7. Scotland In northwest Scotland, United Kingdom, the Loch Maree Group records the LJE with δ13C values ≤ + 13‰ (Baker and Fallick, 1989a). These rocks overlie Archaean basement (Lewisian Complex) that, some 50 km north of the Loch Maree Group outcrop belt, is cut by granitic pegmatites dated at 2480 ± 1 Ma (Corfu et al., 1994) and the Scourie dyke swarm. The latter is sub-divided into four distinct 7 suites, with a bronzite picrite suite dated at 2418+ − 4 Ma and a 3 Ma (Heaman and younger olivine gabbro suite dated at 1992+ −2 Tarney, 1989). On the basis of whole rock geochemistry and petrography, Park (2002) assigned the Scourie dykes underlying

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Fig. 10. Stratigraphic sections and location map of Lomagundi–Jatuli Event-bearing sections, and relevant Palaeoproterozoic sections, from Western Australia. (A) Stratigraphic sections from the Ashburton, Hamersley, Yerrida and Earaheedy Basins after Lindsay and Brasier (2002). See Fig. 4 for the legend. Age references are given in Table 1. References to δ13C are given in Table 2. (B) Map of key Australian Lomagundi–Jatuli Event locality.

the Loch Maree Group to the ‘early quartz–dolerite suite.’ This suite remains to be dated, yet cross-cutting relationships suggest they will be contemporaneous with, or older than, the c. 2.4 Ga bronzite picrite suite (Tarney, 1973; Park, 2002). A detrital zircon study on meta-sedimentary rocks of the Flowerdale Schist in the Loch Maree Group (uncertain stratigraphic relationship to LJE-bearing Loch Maree Group rocks) highlighted two distinct age populations at c. 2200 and 2000 Ma (Whitehouse et al., 1997), including a concordant grain dated at 2474 ± 4 Ma, comparable to the c. 2480 Ma granite pegmatite age from the Lewisian Complex. The Ard Gneiss (granodiorite) cross-cuts the Loch Maree Group and is dated at 1907 ± 3 Ma (Park et al., 2001).

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Fig. 11. Location and stratigraphic sections for South America. See Fig. 4 for the legend. Age references are given in Table 1. References to δ13C are given in Table 2. (A) Map of key South American Lomagundi–Jatuli Event localities. (B) Stratigraphic section from the Minas Supergroup, Brazil, after Bekker et al. (2003b) (C) Stratigraphic section from the San José Greenstone Belt, Uruguay, after Maheshwari et al. (2010).

5.8. Other localities LJE-bearing sedimentary rocks are also located in Norway, Sweden, India and China (see Supplementary Fig. A1 for locations). However, the timing and/or stratigraphic context in those regions is poorly constrained in comparison to the areas discussed in Section 5. In northwest Norway, marbles of the Lofoten–Vesteralen islands have δ13C values ≤ +12.1‰ (Baker and Fallick, 1989b) and are cross-cut by the Lødingen granite dated at 1873 ± 2 Ma (Corfu, 2004), but no maximum age is defined. The Kalix Greenstone Belt in Sweden records δ13C values ≤ +8‰ (Melezhik and Fallick, 2010) in the Upper Formation that is cross-cut by the Jörn Granites dated at 1891 ± 7 Ma (Wilson et al., 1987). The Jhamarkotra Formation, India, has δ13C values between −1.5 and +11.9‰ (Purohit et al., 2010), with the minimum age for the formation provided by a charno-enderbite suite of intrusions dated at 1716 ± 6 Ma (Buick et al., 2006) and a maximum age from a granitoid in the Mewar Gneissic Complex dated at 2440 ± 8 Ma (Wiedenbeck et al., 1996). In China, δ13C values ≤ +5.9‰ (Tang et al., 2010) are found in the Guanmenshan Formation of the North Liaohe Group that is interpreted as part of the Palaeoproterozoic Liabei Terrain on the Sino-Korean Craton (Tang et al., 2010). 6. Discussion Many of the reviewed ages in Section 5 (and Table 1) can be assigned into a pre-, syn- or post-LJE category based upon their stratigraphic context and, in rare examples, where dated horizons co-occur with carbonate rocks having δ13C ≥ 5‰. This approach circumvents many of the difficulties of assigning an age to a specific sedimentary unit, as

discussed in Section 4, and allows the age-categorisation approach as depicted in Fig. 12: the range for pre-LJE dates are shaded green, those for post-LJE dates are shaded blue and the intervening white region indicates the duration permissible for the LJE for a given geographic region. 6.1. Duration of the Lomagundi–Jatuli Event Fig. 12 shows there is a range of maximum, permissible durations for the LJE, from 262 ± 9 Myr based on the South Africa data, to 297 ± 16 Myr based upon the Great Lakes area of North America data, to N 1 Gyr for regions with the least robust age constraints. Fig. 12 also highlights the constraints by which to assess the likelihood that the LJE is a globally synchronous event. When examined region-by-region, there are no occurrences where a post-LJE datum in one region is older than a preLJE datum in another region (Figs. 12, 13). This permits making the assumption made by many studies that the LJE is a globally synchronous signal of positive δ13C values (Baker and Fallick, 1989a; Karhu and Holland, 1996; Bekker et al., 2003a; Melezhik et al., 2005; Maheshwari et al., 2010). In North America, normal δ13C values occur as late as at least 2306 ± 9 (Sturgeon Quartzite; datum 58, Figs. 12, 13) followed by normal δ13C values by c. 2060, with a preferred age at 2056.6 ± 0.8 Ma (2057 ± 1 Ma; datum 30, Figs. 12, 13). These define a maximum duration of 249 ± 9 Myr (2σ; uncertainties added in quadrature) for the LJE. Data assigned to the post-LJE category overlap, within uncertainty, from localities in Finland, Russia, South Africa and Western Australia (Table 1, Fig. 12). The 2058 ± 14 Ma age on detrital zircon from the Chiall Formation, Western Australia (datum 32, Fig. 12) is a

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Fig. 12. Summary compilation of U–Pb (and Re–Os) ages from key Lomagundi–Jatuli Event-bearing sections around the world. References to radiometric dates are shown in Table 1.

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Fig. 13. Schematic diagram of the permissible and measured age constraints for the Lomagundi–Jatuli Event. References to radiometric dates are shown in Table 1. Symbol key is given in Fig. 12.

maximum age and a zircon age of 1891 ± 8 Ma from a tuff bed in the underlying Frere Formation (datum 13, Fig. 12) is a depositional age. This implies that the LJE could have terminated in Fennoscandia and South Africa within at most a few million years of each other. The initiation of the LJE in these areas is less well constrained (Figs. 12, 13). In the Huronian Supergroup the minimum age of the LJE-bearing Gordon Lake Formation is c. 2215 Ma (datum 55, Fig. 12), the age of the cross-cutting Nipissing intrusions. In the Labrador Trough the LJEbearing Portage Formation is older than the cross-cutting Cramolet Lake Sill at 2169 ± 4 Ma (datum 49, Fig. 12) and the Dunphy Formation is older than the Mistamisk Formation (that sits above the overlying Bacchus Formation) dated at 2142 ± 4 Ma (datum 47, Fig. 12). In South America a minimum age for the LJE-bearing Piracicaba and Itabira groups in Brazil is constrained by the 2125 ± 4 Ma date in the Sabará Group (datum 43, Fig. 12) and in Uruguay it is 2146 ± 7 Ma on a rhyolite lava flow overlying the LJE-bearing Carbonate Member (datum 48, Fig. 12). In the four localities listed above, no maximum age constraint b 2.4 Ga currently exists. However in Finland a useful minimum duration for the deposition of LJE-bearing formations can be determined between emplacement of the Laurila Sill, dated at 2221 ± 5 Ma (datum 56 on Figs. 12, 13) and eruption of pyroclastic rocks in the Hirsimaa Formation, dated at 2106 ± 8 Ma (datum 40 on Figs. 12, 13), a period of 128 ± 9.4 Myr (2σ; uncertainties added in quadrature). Ages on sills in the Peräphoja Schist Belt at 2140 ± 11 Ma (datum 46 on Fig. 12) and 2215 Ma (datum 54 on Fig. 12) and in the Kuusamo Schist Belt at 2206 ± 9 Ma (datum 50 on Fig. 12) support this duration. This is a minimum duration for the Peräphoja Schist Belt, and assuming worldwide synchronicity of the LJE, a minimum duration for the LJE excursion. It is noteworthy that the minimum ages listed above for North and South America overlap with the minimum duration seen in the Peräphoja Schist Belt (Fig. 12). 6.2. Global versus diachronous nature of the Lomagundi–Jatuli Event The LJE has been variably interpreted as being either a single, sustained global excursion (e.g. Karhu and Holland, 1996) or consisting of multiple, diachronous excursions (e.g. Melezhik et al., 1999). Given the available age data shown on Fig. 12, the former interpretation is permissible but the latter cannot be dismissed. If the plate configuration shown in Fig. 2 is correct, LJE-bearing sections would have been dispersed at the time of their deposition and given that current age data permit their contemporaneity (Fig. 12) then the LJE was a global event, likely with regional influences superimposed. 6.3. Timing of the Lomagundi–Jatuli Event compared to other key Archaean to Palaeoproterozoic events Understanding the nature of the LJE and its role in the Great Oxidation Event requires consideration of the timing of proxy records of atmospheric oxygen and related environmental records. Huronian

glacial deposits are known from several localities worldwide and wellstudied sections include those from South Africa, Australia and the eponymous Supergroup in Canada. Minimum age constraints for the youngest preserved glaciogenic strata in these localities are given by the 2316 ± 7 Ma age of authigenic pyrite in the Rooihoogte Formation, the 2208 ± 10 Ma age of a sill intruding the Meteorite Bore Member and the 2217 ± 4 Ma Nippising intrusions. Whilst these are minimum age constraints and do not necessarily coincide with the depositional age of the youngest glacial deposit, the age of the LJE as constrained in this paper shows that it is possible that the initiation of the LJE may overlap with that youngest of the Huronian glacial deposits, as supported by various chemo- and litho-stratigraphic correlation schemes (see Hoffman, 2013; Melezhik et al., 2013 for recent reviews). Additionally, the time interval of the LJE coincides with the first appearance of red beds and the disappearance of reducing condition U-ores, post-dates the loss of mass independent fractionation of sulphur, and coincides with a hiatus in iron deposits (Fig. 1). This makes it one of the last, major changes in the record of stable isotope proxies for the Great Oxidation Event (the assumed worldwide organic-C burial episode known as the Shunga event is younger, tending to occur stratigraphically above LJE-bearing rocks; e.g. Melezhik et al., 1999; Kump et al., 2011). Thus, any model attempting to explain the genesis of the LJE must take into account: (i) high δ13C values by at least 2.22 Ga and possible as early as 2.3 Ga; (ii) several of the key Great Oxidation Event proxies pre-date or partially overlap with the initiation of the LJE; (iii) a duration of the LJE of at least 128 ± 9.4 Myr and possibly as long as 249 ± 9 Myr; and (iv) possible contemporaneous termination, within error, on Fennoscandia and South Africa. 7. Targets for future chronological work The termination of the LJE is well constrained in Fennoscandia and South Africa and data also indicate a likely termination of c. 2057 Ma in North America, Siberia and possibly Zimbabwe (Fig. 12). The initiation of the LJE is more poorly constrained (Fig. 13). The Zimbabwe Craton offers the best potential for refining this constraint such that, if the relationship between the Chimbadzi Hill layered complex and LJE-bearing rocks of the Deweras Group could be confirmed, it may extend the duration of the excursion by c. 40 Myr. A sill in the Rukatuntari Formation in Finland's Kuusamo Schist Belt cross-cuts LJE-bearing formations and any age derived from this sill will provide a minimum age for the LJE (Karhu, 1993; Fig. 7B). Fractions of titanite from that sill have been dated by the U–Pb ID-TIMS method at 2078 ± 8 Ma (Silvennoinen, 1991) but whether the titanite is magmatic (Silvennoinen, 1991; Melezhik and Fallick, 2010) or metamorphic in origin (Hanski et al., 2001) remains to be determined thus making the nature of this age ambiguous. In Australia, the Wooly Dolomite contains syn-sedimentary tuffs dated at 2031 ± 15 Ma interbedded with carbonate rocks having normal (≤+2‰) δ13C values (Fig. 10A). This scenario is complicated by a single δ13C value of +6‰ high in the stratigraphy (Lindsay and Brasier, 2002) and resolving how this value relates to the LJE is an important target for future research. Finally, finer-scale resolutions of the time constraints on units that are coeval with the LJE are needed to assess more robustly models advocating the single versus multiple excursion character of the LJE. 8. Conclusions and concluding remarks Stratigraphic sections recording LJE-bearing rocks are known from all continents bar Antarctica. A total of 97 207Pb/206Pb and Re–Os radiometric ages relevant to the LJE have been reviewed with 68 assigned into categories defined as pre-, syn- or post-LJE, with particular attention given to sections in Fennoscandia (Imandra–Varzuga Greenstone Belt, Pechenga Greenstone Belt, Onega Basin, Kuusamo

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Schist Belt, Kiihtelysvaara Area and Peräphoja Belt), North America (Great Lakes area, Labrador Trough and Wyoming), South America (Brazil and Uruguay), Africa (South Africa, Zimbabwe and Gabon), Australia and Scotland. The U–Pb and Re–Os methodologies are the most suitable chronometers for the Precambrian. In order to be confident that a radiometric age provides a maximum, minimum or syn-depositional constraint, published ages need to be explicitly and clearly described as to whether they are interpreted with or without inference with respect to their geological context and how they relate stratigraphically to carbonate rocks with δ13C data. The reviewed chronological data show that a global LJE excursion is permissible, with a maximum permitted range between 2306 ± 9 and 2057 ± 1 Ma and a minimum permitted range between 2221 ± 5 and 2106 ± 8 Ma, yielding a range in duration from a maximum of 249 ± 9 Ma to a minimum of 128 ± 9.4 Myr. A single, sustained C-isotope excursion is permissible, but the possibility of repeated returns to normal C-isotope values cannot be unequivocally ruled out. The permissible timing of a global signal and the probable wide global dispersion of LJE-bearing sections during this time interval, however, suggest that the LJE is a global, rather than diachronous, event. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.earscirev.2013.10.006. Acknowledgements APM, DJC and ARP are supported by NERC grant NE/G00398X/1. Bruce Eglington and David Evans are thanked for useful discussion. We thank A. Bekker and one anonymous reviewer for their reviews and P. Wignall for editorial handling. References Aharon, P., 2005. Redox stratification and anoxia of the early Precambrian oceans: implications for carbon isotope excursions and oxidation events. Precambrian Res. 137, 207–222. Altermann, W., Nelson, D.R., 1998. Sedimentation rates, basin analysis and regional correlations of three Neoarchaean and Paleoproterozoic sub-basins of the Kaapvaal Craton as inferred from precise U–Pb zircon ages from volcaniclastic sediments. Sediment. Geol. 120, 225–256. Amelin, Y.V., Heaman, L.M., Semenov, V.S., 1995. U–Pb geochronology of layered mafic intrusions in the eastern Baltic Shield: implications for the timing and duration of Paleoproterozoic continental rifting. Precambrian Res. 75, 31–46. Anbar, A.D., Duan, Y., Lyons, T.W., Arnold, G.L., Kendall, B., Creaser, R.A., Kaufman, A.J., Gordon, G.W., Garvin, J., Buick, R., 2007. A whiff of oxygen before the Great Oxidation Event. Science 317, 1903–1906. Armstrong, R.A., Wilson, A.H., 2000. A SHRIMP U–Pb study of zircons from the layered sequence of the Great Dyke, Zimbabwe, and a granitoid anatectic dyke. Earth Planet. Sci. Lett. 180, 1–12. Aspler, L.B., Chiarenzelli, J.R., 1998. Two NeoArchaean supercontinents? Evidence from the Paleoproterozoic. Sediment. Geol. 120, 75–104. Baker, A.J., Fallick, A.E., 1989a. Evidence from Lewisian limestones for isotopically heavy carbon in two-thousand-million-year old sea water. Nature 337, 352–354. Baker, A.J., Fallick, A.E., 1989b. Heavy carbon in two-billion year-old marbles from Lofoten-Vesteralen, Norway: implications for the Precambrian carbon cycle. Geochim. Cosmochim. Acta 53, 1111–1115. Barley, M.E., Pickard, A.L., Sylvester, P.J., 1997. Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago. Nature 385, 55–58. Bau, M., Romer, R.L., Lueders, V., Beukes, N.J., 1999. Pb, O, and C isotopes in silicified Mooidraai Dolomite (Transvaal Supergroup, South Africa): implications for the composition of Paleoproterozoic seawater and “dating” the increase of oxygen in the Precambrian atmosphere. Earth Planet. Sci. Lett. 174, 43–57. Bekker, A., Holland, H.D., 2012. Oxygen overshoot and recovery during the early Paleoproterozoic. Earth Planet. Sci. Lett. 317, 295–304. Bekker, A., Kaufman, A.J., 2007. Oxidative forcing of global climate change: a biogeochemical record across the oldest Paleoproterozoic ice age in North America. Earth Planet. Sci. Lett. 258, 486–499. Bekker, A., Kaufman, A.J., Karhu, J.A., Beukes, N.J., Swart, Q.D., Coetzee, L.L., Eriksson, K.A., 2001. Chemostratigraphy of the Paleoproterozoic Duitschland Formation, South Africa: implications for coupled climate change and carbon cycling. Am. J. Sci. 301, 261–285. Bekker, A., Karhu, J.A., Eriksson, K.A., Kaufman, A.J., 2003a. Chemostratigraphy of Paleoproterozoic carbonate successions of the Wyoming Craton: tectonic forcing of biogeochemical change? Precambrian Res. 120, 279–325. Bekker, A., Sial, A.N., Karhu, J.A., Ferreira, V.P., Noce, C.M., Kaufman, A.J., Romano, A.W., Pimentel, M.M., 2003b. Chemostratigraphy of carbonates from the Minas Supergroup, Quadrilatero Ferrifero (Iron Quadrangle), Brazil: a stratigraphic record of early

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