ABSTRACT Large bolide impacts in the Phanerozoic produced global change identifiable in the postimpact sediments. Aside from a few isolated examples, however, evidence of postimpact change associated with Precambrian impacts is sparse. This study used the Neoarchean Paraburdoo spherule layer as a case study to search for impact-induced change in the sediments above the spherule layer. We found possible minor sedimentary changes that may have been due to either a disturbance by bottom currents or changing diagenetic conditions. Contrary to the trends found with several post–Great Oxidation Event large bolide impacts, we found no evidence of shifts in tectonic regime, sediment weathering and deposition, or paleoenvironment induced by the Paraburdoo spherule layer impact, for which the impactor is estimated to have been approximately three times larger than the Cretaceous–Paleogene bolide. This lack of a clear signal of climatic shift may be due to one or more mechanisms. Either the Paraburdoo spherule layer’s deposition in several-hundred-meter-deep water within the Hamersley Basin of Western Australia was too deep to accumulate and record observable changes, or the Neoarchean’s high-CO 2 atmospheric composition acted as a threshold below which the introduction of more impact-produced gases would not have produced the expected climatic and weathering changes. We also report minor traces of elevated iron and arsenic concentrations in the sediments immediately above the Paraburdoo spherule layer, consistent with trends observed above other distal impact deposits, as well as distinctive layers of hematite nodules bracketing the spherule layer. These geochemical changes may record ocean overturn of the Neoarchean stratified water column, which brought slightly oxygenated waters to depth, consistent with the observation of tsunami deposits in shallower impact deposits and/or heating of the global oceans by tens to hundreds of degrees Celsius in the wake of the Paraburdoo spherule layer impact. Either or both of these mechanisms in addition to impact-induced shallow-water ocean evaporation may also have caused a massive die-off of microbes, which also would have produced a postimpact increase in iron and arsenic concentrations.

Abstract The Hamersley province of northwest Western Australia is one of the world's premier iron ore regions. The high-grade iron ore deposits are mostly hosted within banded iron formation (BIF) sequences of the Brockman and Marra Mamba Iron Formations of the Hamersley Group and consist of two types: martite-microplaty hematite containing between 60 and 68 wt percent Fe, and martite-goethite containing between 56 and 63 wt percent Fe. Examples of martite-microplaty hematite include Mount Whaleback, Mount Tom Price, and Paraburdoo and examples of martite-goethite ore deposits include Mining Area C (Area C), Hope Downs, and the Chichester Range. The high-grade martite-microplaty hematite ores, which formed in the Paleoproterozoic, have a three-stage origin. Stage 1 involved the release, from the underlying sedimentary successions, of low (110°C) to high (280°C) temperature, highly saline (20–25.5 wt % NaCl-CaCl 2 equiv; Ca > Na > K) basinal brines that interacted with the underlying Wittenoom Formation and moved upward in normal faults, such as the Southern Batter fault at Mount Tom Price, the 4E fault at Paraburdoo, and the Central and Eastern Footwall faults at Mount Whaleback, into the host BIF. The hypogene fluids migrated laterally within large-scale folds with permeability controlled by shale layers and northwest-trending dolerite dike sets. The BIF was laterally and vertically altered into magnetite-siderite-stilpnomelane and hematite-ankerite ± magnetite assemblages at Mount Tom Price, a hematite-dolomite-chlorite-pyrite assemblage at Paraburdoo, and formed a dolomite-chlorite assemblage in the Mount McRae Shale at Mount Whaleback. Stage 2 involved deeply circulating, low-temperature (<110°C), Na-rich meteoric waters that interacted with evaporites prior to their interaction with the BIF. The descending meteoric waters interacted with the carbonate-altered BIF to produce a martite-microplaty hematite-apatite assemblage prior to supergene alteration. Stage 3, the supergene stage during the Mesozoic to Tertiary, is the final stage in the transformation of BIF to high-grade ore. Shallow supergene fluids interacted with the martite-microplaty hematite-apatite assemblage to form a highly porous high-grade (>63 wt % Fe) martite-microplaty hematite ore. Supergene alteration is likely to have occurred for at least 80 m.y. and close to the present topographic surface. High-pressure (>0.10 wt %) martite-microplaty hematite assemblages can therefore form and may remain concealed beneath BIF, below Proterozoic erosion surfaces. The martite-goethite bedded orebodies resulted from late Mesozoic supergene alteration of BIF. During this process magnetite was oxidized to martite, whereas silicates and carbonates were oxidized and hydrated to goethite or leached without replacement. The controls on the localization of supergene martite-goethite deposits, for example, the Hope Downs, Cloud Break, and Area C deposits include preexisting structures, such as faults, thrusts, and folds. These structures acted as fluid conduits that directed descending supergene fluids into the host BIF. Dolerite dikes and shale layers further focused and controlled fluid flow. High iron grades at the Area C and Hope Downs deposits are associated with synclinal structures where increased supergene fluid flow caused multiple phases of goethite leaching, precipitation, and cementation. Microplaty hematite encompasses a variety of sizes, ranging from 20 to 300 μm, and textures, ranging from platy to tabular. Microplaty hematite is commonly associated with supergene-modified hydrothermal deposits but can also form in the hydration zone of supergene deposits. The phosphorus (P) in supergene and supergene-modified hydrothermal deposits was repeatedly remobilized by both hypogene and/or supergene fluids. The P distribution was controlled by several factors, such as fluid flux in fault zones, permeability of shale layers, and synclinal folds, which resulted in locally high concentrations (>0.10 wt %) of P in the deposits. It is unlikely that a single model for the formation of the martite-microplaty hematite ore deposits can explain all the structural, stratigraphic, hypogene alteration, and ore characteristics at the Mount Whaleback, Mount Tom Price, and Paraburdoo deposits. Continued collaborative research directed at elucidation of a single tectonic history of the Pilbara, based on collection of similar structural and geochemical data sets from these deposits, will advance genetic ore models and aid in exploration for concealed orebodies.