He chronology of the Robe River channel iron deposits, Hamersley Province, Western Australia

Chemical Geology 354 (2013) 150–162

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(U–Th)/He chronology of the Robe River channel iron deposits, Hamersley Province, Western Australia Martin Danišík a,b,⁎, Noreen J. Evans a,c, Erick R. Ramanaidou c, Brad J. McDonald a,c, Celia Mayers c, Brent I.A. McInnes a,c a b c

John de Laeter Centre for Isotope Research, Applied Geology, Curtin University of Technology, GPO Box U1987, Perth, WA 6845, Australia Department of Earth and Ocean Sciences, Faculty of Science and Engineering, The University of Waikato, Private Bag 3105, Hamilton 3240, New Zealand CSIRO Earth Science and Resource Engineering, ARRC, 26 Dick Perry Avenue, WA 6151, Australia

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Article history: Received 13 June 2011 Received in revised form 26 October 2012 Accepted 11 June 2013 Available online 21 June 2013 Editor: K. Mezger Keywords: (U–Th)/He dating Goethite/hematite Channel iron deposits Iron ore genesis Robe River/Mesa J Western Australia

a b s t r a c t Channel iron deposits (CID) supply 40% of Australia's iron ore but their genesis is still the subject of debate. Two well-characterised samples of goethite/hematite CID from a diamond drill core in Mesa J of the Robe River area in Western Australia were dated using (U–Th)/He methods in order to constrain the timing of iron oxide formation and thereby provide a temporal context for CID genesis. (U–Th)/He ages (He ages) range from 25.7 ± 0.6 to 7.0 ± 0.8 Ma and, despite a high degree of scatter, they corroborate relationships expected from the internal ooidal stratigraphy. For individual ooids, the hematitic core is older than or indistinguishable from the age of the surrounding goethitic cortex. The goethitic cortices are, in turn, older than the ferruginised wood fragments recovered from the cementing goethitic matrix. The data suggest the following paragenesis: (i) Hematitic cores in ooids formed in the Early to Middle Miocene as documented by ages of ~ 14.3 ± 3.7 Ma and 18.3 ± 3.5 Ma measured in the shallower (8.2 m deep) and deeper (32.8 m) sample, respectively; (ii) Goethitic cortices of both samples formed in the late Middle to early Late Miocene at 11.6 ± 3.0 Ma; (iii) Wood fragments form a prominent component of the matrix and were ferruginised during the Late Miocene (He ages ranging from 9.4 ± 0.5 to 8.2 ± 0.4 Ma in the deeper core and 8.4 ± 0.9 to 7.0 ± 0.8 Ma in the shallower core). The data suggest that the unique environmental conditions conducive to CID formation existed during the Miocene and that a “typical Robe River CID sequence” likely took 4 to 8 Myr to accumulate. A methodological implication of this study is that it confirms the previous observation of Vasconcelos et al. (2013) suggesting that the temperature utilised for He-extraction from iron oxides has a critical impact on the mobility of parent nuclides. The typical ~ 1000 °C laser heating used for crystalline minerals like apatite or zircon induces loss of U and Th and results in erroneously old ages. Modest extraction temperature (b500 °C), utilising a low-power laser or, preferably, a temperature-controllable resistance furnace is recommended. © 2013 Elsevier B.V. All rights reserved.

1. Introduction (U–Th)/He dating is one of the few tools capable of dating low-temperature geological processes. The vast majority of (U–Th)/ He studies have utilised zircon and apatite, however, given their abundance and high binding affinity with uranium (Duff et al., 2002), secondary iron oxides and oxy-hydroxides are highly perspective candidates for (U–Th)/He dating.

⁎ Corresponding author at: Department of Earth & Oceanic Sciences, The University of Waikato, Private Bag 3105, Hamilton, New Zealand. Tel.: +64 7 838 5163; fax: +64 7 856 0115. E-mail addresses: [email protected] (M. Danišík), [email protected] (N.J. Evans), [email protected] (E.R. Ramanaidou), [email protected] (B.J. McDonald), [email protected] (C. Mayers), [email protected] (B.I.A. McInnes). 0009-2541/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.chemgeo.2013.06.012

The (U–Th)/He dating of iron oxides was first attempted more than 100 years ago (Strutt, 1905, 1910) and its potential as a reliable chronometer was demonstrated two decades ago (Bähr, 1987; Lippolt et al., 1993; Wernicke and Lippolt, 1993; Bähr et al., 1994; Lippolt et al., 1995, 1998). However, recent advances in the understanding of 4He behaviour in iron oxides and oxy-hydroxides (Shuster et al., 2005; Blackburn et al., 2007; Farley and Flowers, 2012) have led to the development of a viable tool for dating those geological processes that cannot be dated by other methods. These include the dating of weathering processes (e.g., Shuster et al., 2005) or ore formation events (e.g., Lippolt et al., 1995; Heim et al., 2006) where iron oxides are often the only datable mineral phases. In this study, we apply (U–Th)/He dating in combination with other analytical methods to investigate the channel iron deposits (CID) of Western Australia. CID are iron-rich fluvial deposits formed in ancient river channels found, with one exception, exclusively in

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Western Australia (Ramanaidou et al., 2003). High iron content (up to 62% Fe), low hardness, amenability to open cut mining and favourable location close to existing infrastructure make CID one of Australia's primary sources of iron ore. Today, CID supply 40% of the iron mined in Australia and industry forecasts indicate that this number will remain constant in the coming decades. Despite their economic importance, models of CID genesis are the subject of debate due to a paucity of geochronological data that could constrain the timing of their formation (MacLeod et al., 1963; Campana et al., 1964; Harms and Morgan, 1964; Butler, 1976; Hall and Kneeshaw, 1990; Morris et al., 1993; Stone, 2005; Heim et al., 2006; Morris and Ramanaidou, 2007; Vasconcelos et al., 2013). This is given by the absence of fossils (other than very abundant wood fragments) and minerals suitable for traditional radiometric dating, as well as by the extreme weathering of the CID. The CID are clearly of Phanerozoic age as the palaeochannels they infill crosscut the Precambrian basement rocks (granitoids, volcanics, metasediments, and banded iron formation; Fig. 1). Morris et al. (1993) and Morris (1994) argued for a Miocene formation age as the palaeochannels incise a so-called Pilbara ferricrete (laterite) palaeosurface of probable Palaeogene age and also fluvial Palaeogene sediments. Estimates based on palynological data could not provide a more accurate age determination and it is currently accepted that filling of the existing palaeochannels could have taken place from Late Eocene to Middle Miocene times (Morris et al., 1993; Morris, 1994; MacPhail and Stone, 2004; Morris and Ramanaidou, 2007; Morris et al., 2007). At the time of writing the only radiometric age constraints were presented by Heim et al. (2006) who applied combined (U–Th)/He and 4He/3He methods to date late-stage authigenic goethite from 10 samples collected from a discontinuous vertical section through the Yandi CID in the Hamersley Province (Western

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Australia). The authors found that the He ages were all Miocene (raw He ages: ~ 14–4 Ma; He ages corrected for diffusive loss: ~ 18– 6 Ma) and progressively decrease with depth, suggesting that precipitation of goethite cement was controlled by water table drawdown during the Neogene (Heim et al., 2006). However, despite having reported ages consistent with the generally accepted Neogene formation age of the CID, the model of Heim et al. (2006) was challenged by Morris et al. (2007), who questioned the sampling strategy and statistical treatment of the data (for details see Morris et al., 2007 and reply by Vasconcelos et al., 2007). Here we aim to (i) derive new age constraints on the formation of CID in the Hamersley Province, (ii) re-evaluate the existing genetic models in the light of the new data, (iii) test the potential and the limitations of the dating technique applied and, where possible, (iv) potentially provide new constraints for the reconstructions of palaeotopography, weathering history and climatic conditions in Western Australia. Samples which represent a typical CID iron ore from the Robe River CID in the Hamersley Province (Figs. 1 and 2) were selected, comprehensively characterised, selectively separated and then dated by (U–Th)/He methods using a range of experimental procedures. 2. Samples and methods For this study, two CID samples (G-82 and G-328) were collected from a diamond drill core (J1136 generously given to CSIRO by Robe River Mining) located at Mesa J in the Robe River Valley (21°44′ 27.23″S 116°15′13.46″E) from vertical depths of 8.2 and 32.8 m below the present surface, respectively (Figs. 1 and 2). Our CID samples formed at low temperature as inferred from hydrogen and oxygen isotopes data on similar CID samples collected in the same

Fig. 1. Upper panel: location (inset) and simplified geological map of the study area. Lower panels: a typical landscape of CID mesas in Western Australia (left and centre); the investigated Mesa J in the Robe River Valley (right).

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Fig. 2. A — Ooidal texture of sample G-328 from the investigated J1136 drill core in Mesa J where three typical components are visible: Hematitic core (1), Goethitic cortex (2) and Goethitic matrix (3). B: Goethitised wood fragment in the matrix. C — SEM image of the wood showing the typical porous structure.

deposit (Mesa J), suggesting formation temperature of 18 ± 3 °C (Yapp personal communication). Macroscopically, both sample specimens represent typical granular type CID ore as described by Morris et al. (1993) (Figs. 2, 3A, and 4A).

They are dominated by red-brown rounded or spherical components (granules), mostly ooids (0.25–2 mm in diameter) with some pisoids (2–5 mm), and less abundant peloids (for terminology see Ramanaidou et al., 2003). All are typified by hematitic cores rimmed

Fig. 3. CID sample G-328 (drill core J1136 328, depth 32.8 m). A: Microphotograph in reflected light displaying the ooidal texture, hematitic nuclei containing some wood fragments, goethite cortices and matrix with goethitised wood fragments. Yellow rectangles labelled B and C show the location of the SEM scans. B and C: X-ray mapping of ooids, cortices and matrix. EDS point analyses on the samples are numbered from 1 to 4 and their alumina, silica and iron oxide content in wt.% are displayed in the inset table D.

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Fig. 4. CID sample G-82 (drill core J1136 82, 8.2 m depth). A: Microphotograph in reflected light displaying the ooidal texture, hematitic nuclei containing some wood fragments (W–Hm), goethite cortices (Co–Go) and matrix (Ma–Go) with goethitised wood fragments (W–Go) with the rectangles showing the location of the SEM scans. B and C: X-ray mapping of ooids, cortices and matrix (back scattered electron image or BSE). EDS point analyses on the samples are numbered from 1 to 4 and their alumina, silica and iron oxide content in wt.% are displayed in the inset table D.

by a narrow (b 0.4 mm) goethite cortex (skin) with a striking metallic lustre. Both samples are massive without obvious bedding; fabric is grain-supported; matrix of fine-grained goethite (sometimes termed as goethitic mud) is porous and contains small voids (b2 mm) without secondary filling, tiny (b 1 mm) fragments of fossilised wood, and irregular unstructured hematitic, goethitic and silicified particles. The deeper sample (G-328) appears to be more heavily leached as inferred from its lighter colour and lower lustre and hardness. It is also less cohesive than the shallower sample (G-82) with a greater tendency to crumble when handled. Prior to dating, petrography, mineralogy, cation substitution and crystallinity of the samples were characterised using petrological analysis, scanning electron microscopy (SEM), energy-dispersive spectroscopy (EDS), X-ray diffractometry (XRD) and Raman spectroscopy. Textural and in situ chemical analyses were respectively completed with an optical microscope ZEISS AXIO Imager.A2m and a scanning electron microscope Phillips XL 40 equipped with an EDS system. XRD patterns were collected with an automated Philips X'Pert MPD diffractometer using CuKα radiation with a post-receiving slit, curved graphite monochromator. Scanning of the sample ranges from 5 to 75° 2θ with 0.03° 2θ steps, counting at 2 s/step. Mineral identification was established using the CSIRO XPLOT® search-match software. For unit-cell determinations, 15 wt.% of 1 μm-sized corundum (α-Al2O3)

was added to some samples as an internal standard to correct for shifts in peak position due to misalignment of the sample during loading and instrumental peak broadening. Aluminium substitution for goethite and hematite was respectively estimated by the method of Li et al. (2006) and Schwertmann et al. (1979). The width of diffraction peaks in an XRD pattern was used to characterise the size and ordering (i.e., ‘crystallinity’) of the samples using the Scherrer equation. Raman spectroscopy analysis was conducted using a Dilor Labram model 1B spectrometer fitted with an 1800 lines/mm dispersion grating, a 150 μm slit and using a Peltier cooled CCD detector at 230 K. Raman spectra were excited using a 14 mW He–Ne laser at 632.82 nm and microscope coupled confocally to a 300 mm focal length spectrograph with detection geometry in back-scattered mode. Focusing of the laser beam on the areas selected for analysis was through a 50 times lens to produce a spot size of approximately 2 μm at the sample surface. Laser power at the sample surface was 1.4 mW, with spectra acquired over a 100 s excitation period. These scanning conditions were selected to prevent any thermal transformation (goethite to hematite) induced in the samples from heating by the laser beam. For (U–Th)/He dating, the samples were coarsely disaggregated in a steel mortar to isolate intact ooids; these were then handpicked under a binocular microscope, ultrasonically cleaned in ethanol, left

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to dry and broken into smaller pieces. When handpicking sample G-328, which shows a slightly more leached appearance, special care was taken to sample the least leached areas. For the first set of samples, hematitic core and goethite cortex shards from different ooids were randomly selected from the clean concentrate. For the second set of analyses, individual ooids were separated into rim and cortex components and then shards from each population were sub-sampled under a binocular microscope. Care was taken to select shards of sufficient size (>150 μm), representing aggregates consisting of small (~ 1 μm) crystals, in order to eliminate the effect of alpha ejection (Farley et al., 1996). In some instances it was possible to pick several shards of sufficient size from the same core and cortex. These were run as multiple aliquots to check the reproducibility of the results. In some cases, ooids were heavily damaged during the crushing and it was not possible to recover both components. Nevertheless, those without pairs were also analysed to generate a more robust dataset. While the striking vitreous lustre made it easy to isolate goethitic cortices, the separation of hematitic cores was more challenging. Care was taken to select homogeneous cores of the same colour, without nuclei. In addition, in each sample it was possible (with difficulty, given the friable nature of the material) to isolate several fragments of ferruginised wood from the matrix. The wood fragments were cut into several shards of sufficient size for analysis, providing multiple aliquots to check reproducibility. Separated core, cortex and wood samples were ultrasonically cleaned in ethanol, dried and loaded into Nb tubes. 4He and 238U with 232Th were measured using isotope-dilution mass spectrometry (quadrupole and ICP-MS, respectively). 147Sm, a possible producer of alpha particles, was not detected in pilot samples and therefore was not included in routine analysis. 4He from the samples was extracted under high vacuum using Nd-YAG laser and analysed on the CSIRO Earth Science and Resource Engineering extraction line at the John de Laeter Centre for Isotope Research in Perth (Australia) on a Pfeiffer Prisma QMS-200 mass spectrometer. Following the traditionally reported extraction temperatures for iron oxides (~ 1150 °C and ~ 1200 °C; Wernicke, 1991; Wernicke and Lippolt, 1993; Shuster et al., 2005), the first set of samples was degassed at ~ 1000 °C for 15 min. However, inspired by the work of other researchers reporting that this high extraction temperatures may cause U loss (P. Vasconcelos, personal communication at TANG3O meeting 2009 in Perth; Vasconcelos et al., 2013) and that secondary iron oxides undergo phase transition at these temperatures (e.g., Boschmann, 1986; Wernicke, 1991; Bähr et al., 1994), extraction temperature was reduced to ~ 500 °C (for 10 min) to circumvent possible changes in the composition of parent isotopes in the samples. Using this heating protocol all our samples were completely degassed in the first heating cycle and no residual gas was observed in the second heating cycle (i.e., during the “re-extraction”; Farley, 2002). This low temperature also served as an independent control for the presence of mineral inclusions (e.g., rutile or zircon) in the samples that might affect the final (U–Th)/ He age. As these minerals can be degassed at temperatures well above 500 °C (e.g., Ehlers and Farley, 2003), their presence would be indicated as a residual 4He signal during the re-extraction. The difference between the results obtained by the two different He extraction procedures is presented in Section 3. Released gas was purified using a ‘cold finger’ cooled with liquid nitrogen and a hot (~350 °C) Ti–Zr getter, spiked with 99.9% pure 3He and introduced into the mass spectrometer next to a cold Ti–Zr getter. 4He/ 3 He ratios, corrected for HD and 3H by monitoring mass 1, were measured by a Channeltron detector operated in static mode. A “re-extract” (Farley, 2002) was run after each sample to verify complete outgassing of the samples. The He gas results were blank corrected by heating empty Nb tubes using the same procedure. Following the 4He measurements, Nb tubes containing the samples were retrieved from the laser cell, spiked with 235U and 230Th

and dissolved in 200 μl of concentrated HCl in Parr bombs heated to 210 °C for 60 h. Each Parr bomb also contained blank and spiked standard solutions. All solutions were analysed for U and Th at TSW Analytical Ltd. (University of Western Australia, Perth) on an Agilent 7500 ICP-MS. For more details on analytical procedures, the reader is referred to Evans et al. (2005). The total analytical uncertainty (TAU) was calculated as a square root of sum of squares of uncertainty on He and weighted uncertainties on U and Th measurements. TAU was typically less than ~ 6% (1 sigma). Given the large size and polycrystalline character of dated samples, (U–Th)/He ages were not corrected for alpha ejection (Farley et al., 1996). The ages were not corrected for diffusive loss (Shuster et al., 2004). 3. Results 3.1. Sample characterisation: Microscopy, SEM, EDS, XRD and Raman spectroscopy The hand specimen of sample G-328 (32.8 m) shows an ooidal texture with hematitic cores (red), goethitic cortices (hard vitreous bluish-grey) and a goethitic matrix with fossilised wood fragments (Figs. 2, 3 and 4). The shallower G-82 sample has lighter colour and lower lustre than the deeper sample (G-328). The petrological study of the two samples under reflected and transmitted light clearly depicts the ooidal texture typical of a granular CID (as defined by Ramanaidou et al., 2003). The vast majority of the granules range from 0.25 to 2 mm in size and comprise a hematitic nuclei surrounded by a goethitic cortex, cemented by a goethitic matrix containing wood fragments. Under reflected light, the hematitic nuclei (cores) can be simple or complex, compact or porous, with or without hematitised wood fragments. Two selected areas in each sample were chosen for the SEM and EDS examination (Figs. 3B–D and 4B–D). The shallower sample revealed that hematitic cores have an alumina content ranging from 2.5 to 3.6% and are lower in alumina than their equivalent cortices (>4%). The alumina and silica values also show that the Al is not only in the structure of the iron oxides but also intimately associated with kaolinite (Si2Al2O5 (OH)4). The deeper sample (Fig. 4) presented a similar ooidal texture and alumina and silica values with a similar pattern to that in the shallower sample, however, the alumina content is generally lower. These chemical variations between various nuclei and cortices are consistent with Robe River valley CID geochemical data published by Ramanaidou et al. (2003). Thus, petrologically, both investigated samples represent a typical CID occurring in the Robe River valley. The XRD patterns (with added corundum as an internal standard) for the two samples show that the hematite and goethite fitted peak positions (Fig. 5A) are respectively centred at 3.682 and 4.183 Å (Fig. 5A, B), characteristic of a low level of Fe–Al substitution (b2 wt.% of Al2O3). In addition, unit-cell dimensions for hematite and goethite is very similar to those of pure hematite, indicating very little or no replacement of Fe by Al in the hematite structure. The width at half height for the (110) of goethite and the (012) of hematite are very similar in both samples and show relatively well-crystallised minerals (Fig. 5B). Wavenumber values (in cm−1) for excitation peaks of the goethite occurring in the cortex and matrix (including wood fragments) and hematite (nucleus of the ooids) (Fig. 6) are consistent with the Raman spectra reported in the literature (de Faria et al., 1997; Bersani et al., 1999; Table 7 in Ramanaidou et al., 2008). The spectra suggest very little Al substitution in the cortex and matrix of both samples. The presence of a Raman peak at 245 cm−1 is characteristic of good crystallinity and low Al substitution (Schwertmann and Taylor, 1989; Ramanaidou et al., 2008). Broadening of the main phonon lines for goethite also reflect its less well-ordered nature as

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Fig. 5. A — Bulk XRD patterns for the two samples (G-328 and G-82) showing hematite, goethite and the added corundum (internal standard). Patterns are vertically offset for clarity. B — Hematite and goethite fitted peak positions respectively 3.682 and 4.183 Å characteristic of low level of substitution. The width at half height for the (110) of goethite and the (012) of the hematite are very similar in both samples and show relatively well crystallised minerals.

shown in the goethitic cortex. The Raman spectrum background of the wood matrix goethite in sample G-82 is high because of the resin that produces fluorescence effect. In summary, XRD and Raman spectroscopy results show that both goethite cortex and matrix, as well as the hematite nucleus, contain very little Al (less than 2 wt.%) and are relatively well crystallised minerals, which justify their suitability for (U–Th)/He geochronology (Shuster et al., 2005). 3.2. (U–Th)/He results (U–Th)/He results are summarised in Table 1. He ages range from 25.7 ± 0.6 to 2.2 ± 0.3 Ma, show significant dispersion and, at first glance, apparently lack a systematic pattern. He contents vary from 0.01 to 0.35 ncc at STP and Th/U ratios vary from ~ 0.01 to ~ 24. He ages measured on multiple aliquots differ by ~ 2 to 25%, however, in most of the cases 2 out of 3 replicates reproduce within 1 sigma error and one replicate is a ‘flyer’ (Table 1). We confirmed previous observations (Vasconcelos et al., personal communication at TANG3O meeting 2009; Vasconcelos et al., 2013) that He extraction temperature has a significant impact on the Th/U ratios and on the ages. Fig. 7 illustrates a comparison of Th/U ratios for G-82, gas extracted at high (~1000 °C) and low (~500 °C)

temperature, where a shift of the Th/U ratio towards lower values for the low extraction temperature is obvious. This indicates that one or both parent isotopes (perhaps preferentially U) were driven off at high temperature in the laser chamber. This undesired opening of the isotopic system must have led to erroneous results (more scattered and potentially erroneously old ages; Fig. 7) so the cortex and core ages obtained during high temperature gas extraction were excluded from further interpretation. The wood fragments from the G-82 shallow core were gas extracted at ~ 1000 °C and paucity of sample precluded a repeat analysis. The structure of the wood is naturally porous (Fig. 2C) and helium retentivity may be less robust than in more crystalline iron oxides. Given these uncertainties, the matrix wood analysis for the shallow core is potentially less reliable. Regardless, the reproducibility of the few ages obtained argues that the data should not be entirely dismissed (Table 1). For the remainder of the data the following four relationships are noted (see also Table 1; Figs. 8 and 9): (i) All He ages from the cortex–core pairs reflects their internal stratigraphic relationship. The He ages of the cores are older than those of the corresponding cortices (25.7 ± 0.6 to 10.5 ± 0.4 Ma and 18.9 ± 1.2 to 9.6 ± 0.2 Ma, respectively; Fig. 9).

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Fig. 6. Raman spectra of goethite and hematite of samples G-82 (A — upper panel) and G-328 (B — lower panel). The cortex and matrix goethite contains very little aluminium substitution as shown using the equation: y = 388 − 3.86 ∗ x − 0.65 ∗ x2 + 0.51 ∗ x3 − 0.001 ∗ x4 (data from Ramanaidou et al., 1996). Similarly the nucleus hematite is Al-poor as shown by the lack of a strong peak at 662 cm−1, well related to Al content (data from Ramanaidou et al., 1996). Note that the background of the wood matrix goethite in G-82 is high due to the fluorescence effect.

(ii) The wood fragment from the matrix of the deeper sample (G-328) yielded ages from 9.4 ± 0.5 to 8.2 ± 0.4 Ma (n = 3; central age: 8.8 ± 0.6 Ma; Fig. 8), clearly younger than He ages of the cores and cortices from that sample. The ages of wood from the matrix of the shallower core (G-82: 8.4 ± 0.9 to 7.0 ± 0.8 Ma; n = 5; central age: 7.5 ± 0.5 Ma) are also consistently younger than the cores and cortices from this depth. This age relationship reflects internal stratigraphic relationships, where the matrix should be the last phase of the CID to form. (iii) There is a difference between the samples in terms of He ages of ooid cores. The near surface sample G-82 yielded core He ages ranging from 22.3 ± 1.0 to 9.4 ± 0.3 Ma with the central age of 14.3 ± 3.7 Ma (n = 14), which are younger than He ages yielded by the deep sample G-328 (ages ranging from 25.7 ± 0.6 to 13.0 ± 0.4 Ma with a central He age of 18.3 ± 3.5 Ma; n = 13; Fig. 8). (iv) He ages of cortices from both samples are indistinguishable (G-82: 11.9 ± 1.8 Ma and G-328: 11.6 ± 3.0 Ma; Fig. 8).

4. Interpretation and discussion 4.1. Scatter of (U–Th)/He ages Precision of He ages is in some cases as high as 25%, which is much poorer than the reproducibility of He ages typically obtained on magmatic minerals such as zircon or apatite that are routinely dated by the (U–Th)/He method. It is therefore critical to assess the reliability

of the data by systematically reviewing all the parameters involved in the measurements. These include the following. 4.1.1. Poorly understood 4He retentivity and closure temperature The early results of 4He diffusion experiments showed that hematite and goethite quantitatively retain radiogenic He (Lippolt et al., 1993; Bähr et al., 1994; Wernicke and Lippolt, 1994a, 1994b) and allowed the 4He closure temperatures (Tc) to be determined for specific textural varieties of these minerals. The reported Tc's are 219 °C for specular (hydrothermal) hematite (grain diameter >5 mm; Lippolt et al., 1993) and 90–160 °C for botryoidal (supergene) hematite consisting of radial micrometre-sized crystals (Bähr et al., 1994). Recent study by Farley and Flowers (2012) reported Tc's of b0 to ~180 °C (depending on size of diffusion domains) for hydrothermal hematite. The studied supergene hematite, however, cannot be described as botryoidal and its crystal structure is potentially different from the sample analysed by Bähr et al. (1994). Since the He retentivity depends on crystallite shape, size, and packing density (Shuster et al., 2005; Farley and Flowers, 2012), assuming this Tc in the current work would not be appropriate. Bähr et al. (1994) also pointed out that a natural He loss of only 5% will lower the apparent calculated closure temperatures to only 40–110 °C, which would result in younger ages than expected. Although this effect can be quantified and corrected using the 4He/3He method (Shuster et al., 2004), this approach was not used in this study and thus the ages could be biased towards the younger values by an unknown factor. For supergene goethite, Shuster et al. (2005) recommended not using Tc for characterising He retentivity and showed that 86–90% of radiogenic 4He was retained. The 4He diffusive-loss corrections,

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Table 1 (U–Th)/He resultsa. Sample code

Pelletoid part

Extraction temperature ~1100 °C G82-C1 Hematitic core -C2 Hematitic core -C3 Hematitic core -C4 Hematitic core -C5 Hematitic core -C6 Hematitic core -C7a Hematitic core -C7b Hematitic core -C7c Hematitic core -C7d Hematitic core G82-R1 Goethite cortex -R2 Goethite cortex -R3 Goethite cortex -R4 Goethite cortex -R5 Goethite cortex -R6 Goethite cortex -R7a Goethite cortex -R7b Goethite cortex -R7c Goethite cortex -R8a Goethite cortex -R8b Goethite cortex G328-C1 Hematitic core -C2 Hematitic core -C3 Hematitic core -C4 Hematitic core -C5a Hematitic core -C5b Hematitic core -C5c Hematitic core -C6 Hematitic core -C7a Hematitic core -C7b Hematitic core -C7c Hematitic core G328-R1 Goethite cortex -R2 Goethite cortex -R3 Goethite cortex -R4 Goethite cortex -R5 Goethite cortex -R6 Goethite cortex -R7a Goethite cortex -R7b Goethite cortex -R7c Goethite cortex -R8a Goethite cortex -R8b Goethite cortex -R8c Goethite cortex Extraction temperature ~500 °C Core–cortex pairs G82-R11a Goethite cortex -R11b Goethite cortex -R11c Goethite cortex -C11a Hematitic core -C11b Hematitic core -C11c Hematitic core -R12a Goethite cortex -R12b Goethite cortex -R12c Goethite cortex -C12a Hematitic core -C12b Hematitic core n/a Goethite cortex -C13a Hematitic core -C13b Hematitic core -C13c Hematitic core -C14 Hematitic core -R14 Goethite cortex -C15 Hematitic core -R15 Goethite cortex -C16 Hematitic core -R16 Goethite cortex -C17 Hematitic core -R17 Goethite cortex n/a Hematitic core -R18 Goethite cortex -C20 Hematitic core -R20 Goethite cortex

Th (ng)

Th err. (%)

U (ng)

U err. (%)

He (ncc)

He err. (%)

TAU (%)

Th/U

Age (Ma)

±1σ (Ma)

0.263 0.267 0.126 0.137 0.256 0.233 0.089 0.223 0.354 0.313 0.178 0.221 0.328 0.221 0.180 0.249 0.214 0.129 0.183 0.101 0.196 0.275 0.239 0.199 0.226 0.258 0.176 0.210 0.148 0.152 0.144 0.091 0.110 0.212 0.098 0.115 0.128 0.345 0.108 0.109 0.130 0.137 0.123 0.155

2.6 2.8 2.8 2.6 2.7 2.7 2.2 2.0 2.0 1.9 2.5 3.2 2.6 2.8 2.6 2.7 2.5 3.2 3.7 2.3 3.2 3.5 2.6 2.7 2.6 3.0 3.2 3.3 4.3 2.3 2.1 3.4 2.6 2.9 2.6 2.8 2.7 1.9 3.4 2.8 3.0 3.8 2.7 2.3

0.025 0.034 0.005 0.017 0.033 0.025 0.005 0.024 0.033 0.037 0.028 0.049 0.052 0.042 0.046 0.031 0.021 0.009 0.015 0.021 0.055 0.041 0.060 0.041 0.070 0.047 0.047 0.034 0.035 0.063 0.043 0.034 0.054 0.086 0.051 0.052 0.074 0.126 0.042 0.073 0.043 0.031 0.042 0.058

1.5 1.5 3.1 2.0 2.0 1.9 2.2 2.6 2.5 2.3 1.9 1.6 1.5 1.9 1.9 2.1 2.2 2.6 2.5 2.3 2.3 1.7 3.8 1.9 3.2 2.5 2.7 2.5 2.7 2.3 2.3 2.6 1.7 2.2 1.5 2.0 1.9 2.3 2.7 3.0 2.4 3.6 2.5 2.3

0.145 0.174 0.076 0.080 0.138 0.133 0.052 0.134 0.200 0.177 0.132 0.186 0.279 0.176 0.178 0.195 0.126 0.081 0.114 0.081 0.184 0.208 0.240 0.144 0.223 0.115 0.099 0.125 0.124 0.180 0.126 0.111 0.138 0.207 0.128 0.141 0.144 0.357 0.152 0.182 0.220 0.156 0.111 0.197

21.4 21.3 5.3 5.0 2.9 3.0 5.8 2.2 3.0 1.7 4.5 4.8 1.8 2.3 2.8 3.0 2.4 3.7 2.6 3.7 1.6 5.8 2.5 2.8 2.7 3.0 3.0 2.6 2.6 3.1 2.4 2.7 10.1 4.3 14.1 2.8 2.8 2.8 3.3 3.0 2.7 3.2 2.7 2.8

21.5 21.3 5.8 5.3 3.5 3.6 6.1 2.8 3.4 2.3 4.9 5.2 2.4 2.9 3.2 3.6 3.0 4.5 3.9 4.1 2.5 6.2 3.4 3.2 3.4 3.6 3.7 3.4 3.6 3.5 2.9 3.4 10.2 4.7 14.1 3.3 3.2 3.2 3.9 3.8 3.3 4.1 3.3 3.3

10.35 7.91 24.41 7.79 7.76 9.12 18.68 9.19 10.61 8.46 6.26 4.53 6.20 5.17 3.86 8.00 9.99 14.01 12.11 4.78 3.55 6.67 3.97 4.86 3.20 5.42 3.68 6.07 4.17 2.41 3.28 2.69 2.03 2.46 1.91 2.17 1.73 2.71 2.54 1.47 3.01 4.37 2.94 2.67

13.7 14.8 17.9 13.2 12.2 13.6 16.6 14.3 14.1 13.2 15.5 15.2 17.7 15.3 16.5 17.9 14.4 16.8 16.1 14.8 14.9 16.2 17.0 13.5 14.9 8.8 9.2 12.3 14.5 15.0 13.4 16.5 14.2 12.5 14.2 14.6 11.4 14.1 18.5 15.2 24.5 20.2 12.9 17.2

2.9 3.2 1.0 0.7 0.4 0.5 1.0 0.4 0.5 0.3 0.8 0.8 0.4 0.4 0.5 0.6 0.4 0.8 0.6 0.6 0.4 1.0 0.6 0.4 0.5 0.3 0.3 0.4 0.5 0.5 0.4 0.6 1.4 0.6 2.0 0.5 0.4 0.5 0.7 0.6 0.8 0.8 0.4 0.6

0.074 0.144 0.117 0.121 0.122 0.071 0.076 0.043 0.067 0.086 0.052

0.1 5.9 5.9 5.9 5.9 6.0 6.0 6.4 6.3 0.1 6.1

0.020 0.077 0.056 0.019 0.012 0.005 0.031 0.014 0.026 0.010 0.007

0.0 2.2 2.1 2.3 2.1 2.3 2.4 2.3 3.2 0.0 2.4

0.066 0.131 0.122 0.061 0.056 0.037 0.078 0.031 0.056 0.038 0.038

1.7 1.0 1.0 1.8 1.9 2.9 1.4 3.4 2.0 2.8 2.9

5.0 3.9 3.9 5.1 5.5 6.1 4.4 5.6 5.0 5.9 5.8

3.60 1.85 2.06 6.25 10.49 14.74 2.43 2.99 2.53 8.88 7.42

14.3 9.6 12.0 10.5 11.4 14.0 13.0 10.4 11.0 10.5 16.1

0.5 0.2 0.3 0.4 0.5 0.8 0.4 0.5 0.4 0.5 0.8

0.142 0.196 0.182 0.136 0.083 0.123 0.062 0.053 0.045 0.056 0.092

0.1 5.9 6.0 5.9 6.1 6.2 6.2 6.1 6.1 3.7 2.1

0.061 0.037 0.038 0.026 0.035 0.018 0.035 0.006 0.025 0.012 0.044

0.0 2.2 2.1 2.2 2.6 2.6 2.4 2.7 2.1 1.9 1.8

0.108 0.182 0.163 0.110 0.075 0.083 0.061 0.050 0.056 0.057 0.095

1.1 0.8 0.8 0.9 1.1 1.0 1.3 1.6 1.4 4.5 3.0

4.1 4.7 4.7 4.7 4.3 5.2 4.2 5.5 4.1 5.4 3.6

2.31 5.32 4.77 3.00 3.38 3.37 3.42 1.60 1.16 4.77 2.10

9.4 18.0 16.5 15.7 11.2 14.6 10.2 22.3 13.0 19.0 11.9

0.3 0.6 0.6 0.6 0.3 0.6 0.3 1.0 0.4 1.0 0.4

0.129 0.097 0.119

1.9 2.4 1.9

0.035 0.011 0.057

1.9 1.9 1.7

0.130 0.050 0.121

2.4 5.1 2.5

3.1 5.5 3.1

3.70 8.79 2.07

16.5 12.2 11.6

0.5 0.7 0.3

(continued on next page)

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Table 1 (continued) Sample code

Pelletoid part

Extraction temperature ~500 °C Core–cortex pairs -C21 Hematitic core -R21 Goethite cortex G328-C11a Hematitic core -C11b Hematitic core -R11a Goethite cortex -R11b Goethite cortex -C12a Hematitic core -C12b Hematitic core -C12c Hematitic core -R12a Goethite cortex -R12b Goethite cortex -C13 Hematitic core -R13 Goethite cortex -C14 Hematitic core -R14 Goethite cortex -C15 Hematitic core -R15 Goethite cortex -C16 Hematitic core -R16 Goethite cortex -C17 Hematitic core -R17 Goethite cortex -C18 Hematitic core n/a Goethite cortex -C19 Hematitic core n/a Goethite cortex -C20 Hematitic core -R20 Goethite cortex Ferruginised wood from the matrix G82b W1a Matrix wood W1b Matrix wood W1c Matrix wood W1d Matrix wood W1e Matrix wood G328 W3a Matrix wood W3b Matrix wood W3c Matrix wood

Th (ng)

Th err. (%)

U (ng)

U err. (%)

He (ncc)

He err. (%)

TAU (%)

0.044 0.045 0.201 0.153 0.109 0.133 0.221 0.132 0.138 0.066 0.083 0.027 0.023 0.043 0.040 0.047 0.063 0.056 0.098 0.054 0.024 0.157

2.2 2.5 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 6.4 6.4 6.2 6.0 6.1 6.1 2.2 2.0 2.2 2.8 2.5

0.007 0.018 0.059 0.044 0.036 0.039 0.079 0.036 0.041 0.041 0.055 0.014 0.020 0.033 0.025 0.029 0.045 0.013 0.029 0.014 0.012 0.050

1.6 1.7 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 2.5 2.2 2.4 2.1 2.2 2.4 2.0 1.7 1.7 1.6 1.6

0.027 0.042 0.204 0.151 0.091 0.111 0.207 0.135 0.173 0.056 0.079 0.037 0.036 0.122 0.049 0.126 0.085 0.060 0.083 0.065 0.041 0.216

8.8 5.9 0.7 0.9 1.3 1.1 1.0 1.2 1.0 2.2 1.7 2.1 2.1 0.8 1.6 0.8 1.0 4.4 3.3 4.1 6.0 1.8

9.0 6.3 4.6 4.4 4.3 4.5 4.2 4.5 4.8 4.3 4.0 4.6 4.1 3.7 4.0 3.7 3.8 4.9 3.8 4.5 6.4 2.7

0.039

3.3

0.033

2.0

0.098

2.9

0.041 0.056

3.7 2.1

0.030 0.043

1.8 1.6

0.091 0.068

0.094 0.141 0.158 0.117 0.186 0.045 0.047 0.027

2.0 2.4 3.2 2.7 2.7 1.8 6.6 2.4

0.007 0.012 0.017 0.017 0.013 0.047 0.053 0.053

2.3 2.3 2.5 2.0 2.0 1.5 1.6 1.6

0.030 0.041 0.048 0.038 0.054 0.067 0.067 0.060

Th/U

Age (Ma)

±1σ (Ma)

6.64 2.44 3.65 3.30 1.50 2.77 1.60 1.58 1.39 1.96 1.29 5.23 2.36 6.71 1.77 8.69 1.80 4.43 3.33 3.84 1.96 3.12

13.3 11.9 15.8 15.5 12.1 12.9 13.0 16.6 19.2 8.2 8.7 15.3 11.6 23.5 11.7 25.7 11.7 19.0 13.1 20.2 18.9 20.5

1.2 0.7 0.5 0.5 0.4 0.4 0.4 0.5 0.6 0.3 0.2 0.5 0.4 0.6 0.3 0.6 0.3 0.9 0.5 0.9 1.2 0.5

3.8

1.17

19.1

0.6

3.1 3.9

3.9 4.3

1.37 1.30

18.9 10.1

0.7 0.4

10.0 7.3 6.3 10.5 7.4 4.6 4.6 4.9

10.1 7.6 6.7 10.7 7.7 4.8 4.9 5.1

13.20 11.23 9.06 6.95 14.27 0.95 0.89 0.51

8.4 7.4 7.2 7.0 7.8 9.4 8.7 8.2

0.9 0.6 0.5 0.8 0.6 0.5 0.4 0.4

Values written in bold represent measured ages. a Th — 232Th; U — 238U; He — 4He at STP; TAU — total analytical uncertainty; Age — raw (U–Th)/He age uncorrected for alpha recoil (see text for explanation). b Note that this matrix wood was gas extracted at 1000 °C so the ages represent maxima.

which can be quantified using the 4He/3He method, could increase the He ages in the present samples by up to ~ 15%.

Fig. 7. Impact of extraction temperature on (U–Th)/He age and isotopic composition of U and Th demonstrated on sample G-82. Extraction temperature of 1000 °C likely mobilised Th and possibly U (note the difference in Th/U ratios), which resulted in erroneously high (U–Th)/He ages.

4.1.2. Relatively low precision and indeterminate accuracy Given the lack of a representative age standard, it is difficult to determine the accuracy and precision of (U–Th)/He dating of secondary iron oxides. Lippolt et al. (1995) validated geological meaningfulness of He ages obtained on specular (hydrothermal) hematite by using K/Ar ages of coexisting adularia and biotite as independent benchmarks. However, similar cross-validation tests performed on supergene goethite dated by (U–Th)/He and coexisting cryptomelane dated by 40Ar/39Ar were less successful (Vasconcelos et al., 1994; Shuster et al., 2005). Consequently, the accuracy and precision of the (U–Th)/He method on supergene minerals are unknown. In order to provide some indication of precision, we performed replicate analyses of a fresh, homogeneous, non-leached, high quality botryoidal goethite from Arkaroola (South Australia). The nature of this sample is distinct from those analysed from the Robe River Valley and it is U-dominated (Th/U ratio 0.001 c.f. ratios typically >1 at Robe River). However, Arkaroola analyses (n = 8) performed over the period of this study yielded a mean (U–Th)/He age of 20.8 ± 3.0 Ma, indicating that the precision of ~20% (1σ) obtained on real geological samples is low (cf. Durango apatite, analysed routinely with a precision better than 5%; n = 252, CSIRO lab). While this precision would perhaps improve with a larger number of analyses, this result also suggests that even worse precision should not be surprising in significantly weathered samples, such as the Robe River CID. There is no “accepted” age for the

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159

Fig. 8. (U–Th)/He results displayed in histograms (left panels) and radial plots (right panels). Note the difference in He age between cores of the ooids from each depth, fairly similar age of cortices in both the 8.2 m and 32.8 m core and distinctly younger age of the wood fragments in both.

Arkaroola goethite so we were not able to determine the accuracy of the methods. 4.1.3. Complex character of the dated samples As illustrated previously (Ramanaidou et al., 2003; Morris and Ramanaidou, 2007) and confirmed by microscopy, both hematitic cores and goethite cortices often consist of a series of layers, creating complex laminated or onion-skin textures (Figs. 3 and 4). The layers represent distinct growth bands, which precipitated and/or possibly ecrystallised at different times and should thus record different (U–Th)/He ages. It was not possible to micro-sample individual bands and separated shards used for analysis may, therefore, represent randomly sampled cross-sections consisting of several layers. Measured He ages may represent an average age for each dated shard, which might deviate somewhat from the true average He age of the whole

cortex/core, thus adding additional uncertainty into the measured age dataset. Moreover, inclusions and other impurities containing U, Th or He, that were undetected during the sample preparation, may have contributed to the observed scatter.

4.1.4. Potential of U and Th loss during He extraction As indicated in Section 3, elevated extraction temperatures (1000 °C) had significant impact on the behaviour of U and Th in the samples. Lower helium extraction temperatures were applied (500 °C) leading to less vaporisation of parent isotopes and more reproducible results. There are, however, no guarantees that the 500 °C temperature is sufficiently low and, if not, the ages will be still be biased towards older values. Unfortunately, the Nd-YAG laser used as a heat source did not couple with the Nb tubes at temperatures of b 500 °C so it was not possible to test this

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Fig. 9. Comparison of (U–Th)/He ages of core–cortex pairs. Overall trend corroborates internal stratigraphic relationships where cores formed before cortices and thus should yield older He ages.

theory. In future experimental studies, an externally heated, low-blank resistance furnace with a well-controlled temperature will be utilised. When all these factors are considered, the observed scatter in He ages is understandable and, perhaps internally consistent and occasionally reproducible ages are somewhat surprising (e.g., G-328, C11 and R11 and deep core matrix wood analyses). However, despite the scatter and limited resolution of the method, the data reveals systematic relationships and provides constraints on the genesis of CID, which will be discussed in the following section. 4.2. Implications for CID genesis Given the lack of low-temperature thermal events during the Miocene period in Western Australia (Kohn et al., 2002) and provided that the (U–Th)/He system in dated samples remained closed (i.e., U, Th and 4He have been immobile since mineral formation), the He ages most likely represent the time of mineral formation and are interpreted as formation ages. The results of the He dating summarised in Section 3 suggest the following succession of ore forming processes: 1. Formation of ooid hematitic cores of the deeper sample (G-328) at 18.3 ± 3.5 Ma; 2. Formation of ooid hematitic cores of the shallower sample (G-82) at 14.3 ± 3.7 Ma; 3. Formation of goethitic cortex at 11.6 ± 3.0 Ma, which coated the pre-existing hematitic cores of both samples; 4. Ferruginisation of the wood fragments in the deeper matrix at 8.8 ± 0.6 Ma and in the shallower matrix at 7.5 ± 0.5 Ma. It should be emphasised that relative timing of the processes is the key information here, as there is some uncertainty associated with the absolute ages as explained above. The existing genetic model for CID formation can be constrained by the new He data as follows: The ages of the cores imply that ooids of dated samples formed in the Early to Middle Miocene. The inverse age–depth relationship between the samples may suggest that ooids represent components derived either from two sources of different age or from the same regolith profile which became inverted during the erosion, transport and re-sedimentation. The deeper (older He ages) and shallower (younger He ages) ooids thus represent the older, near-surface and younger, deeper regolith levels, respectively. The formation of the goethitic cortex in the late Middle to early Late Miocene, in contrast, appears to be a fast process lasting maximally a few million years as inferred from indistinguishable He ages. The location of goethitic cortex formation is somewhat arguable.

Whereas Morris and Ramanaidou (2007) favoured formation of the goethitic cortices during colluvium stage before accumulation of the ooids in the river channels, Heim et al. (2006) proposed that goethitic cortex formed after the channel aggradation and cemented the cores. Heim et al. (2006) also found a downward younging trend in He ages measured over ~ 30 m deep discontinuous profile. These authors argued that the process of iron cementation occurred at the groundwater–atmosphere interface and was driven by water table drawdown during the Neogene. Although we cannot discern from our data which model is correct, we can say that the Robe River CID data do not show the trend observed by Heim et al. (2006). The two cortex samples, with a vertical distance of 24 m, collected in the same borehole do not show any variation in He age. Therefore, keeping in mind the limited number of samples studied, we tentatively suggest that the downward-progressing cementation model driven by water table drawdown cannot apply to the locality investigated in this study. The last process recorded by He data seems to be the ferruginisation of wood fragments, in which the organic wood tissues were replaced by mobilised ferrous iron. Although the mechanism of this process is not well understood (Morris et al., 1993, 2007; Morris and Ramanaidou, 2007), it most likely occurred in the Late Miocene (~10–7 Ma). We would like to emphasise that this conclusion is valid only if the ages record the time of goethite/hematite replacement, and that they were not a result of less robust helium retentivity and/or ejection of alpha particles in the porous and less crystalline wood samples. Nevertheless, this result clearly shows that river aggradation must have been completed in the Late Miocene. When an “average ooid” age was generated for the samples from each depth, the shallow sample yielded a 13.0 Ma age while the deeper sample formed 15.0 Ma. This implies that over the 24 m interval of core, the ooids took about 2 Myr to form, assuming a continuous sequence of CID accumulation. This is consistent with previous modelling and suggests that a “typical” 100 m sequence of Robe CID takes 4–8 Myr to accumulate. This needs to be confirmed on a wider range of core samples from different depths. 4.3. Implications for geomorphological, climatic and weathering history of WA One of the objectives of this study was to provide time constraints on processes and conditions that directly controlled CID formation. These include palaeo-geomorphological setting, climatic conditions and erosional history. Although our data corroborate existing

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estimates, the relatively high scatter in the dataset, limited resolution of the (U–Th)/He dating method and the limited number of cores analysed precluded us from refining the timing of these processes further. The meandering nature and width of the river channels indicate that a mature, low relief landscape with flat topography and low gradients existed in Hamersley in the past. The age of the landscape is believed to be Late Eocene to Middle Miocene as inferred from palynological data at Yandi and relative dating of landforms (Morris et al., 1993; Morris, 1994; MacPhail and Stone, 2004; Morris et al., 2007). The He ages of the ooids that were later accumulated in the channels indicate that the mature landscape with developed regolith must have existed in the Early Miocene, which is consistent with these estimates. It is believed that CID genesis was controlled by unique climatic conditions. Although these are difficult to establish, it is noteworthy that the age estimates obtained using different approaches converge. Morris and Ramanaidou (2007) proposed a tropical to subtropical climate without significant seasonal changes using the presence of angiosperm tropical woodland species and lack of tree rings in fossil wood as arguments. Rainfall was generally gentle and/or catchments were protected by vegetation as evidenced by fine-grained bedrock clastics, although intermittent torrential rains or storms producing coarse-grained bedrock clastics occasionally occurred. These authors used different sources of information including palynology (MacPhail and Stone, 2004), palaeontology (McGowran and Li, 1998), sea-level eustatic curves (Haq et al., 1987), palaeomagnetic data from regolith (Pillans, 2002) and palaeogeography (Veevers, 2001) and concluded that these unique climatic conditions favourable for the CID formation existed in the Hamersley Province since the middle Early Miocene with a climax in the Middle Miocene (so called Miocene climatic optimum). Similarly, Heim et al. (2006) argued that channel aggradation marking a transition toward arid conditions started after the Early Oligocene based on (U–Th)/He data. The channel aggradation was, according to the authors, a fast process creating spaces prone to groundwater accumulation and subsurface flow. Secondary goethite was then precipitated at the groundwater–atmosphere interface propagating downwards throughout the Neogene. Our data fit well the Miocene climatic optimum for CID formation proposed by Morris and Ramanaidou (2007), and although they do not confirm the age–depth trends found by Heim et al. (2006) (see Section 4.2), the absolute He ages of both studies are in accord. Further, our data are conspicuously similar to the results of Pidgeon et al. (2004) and Thern et al. (2011), who reported Late Miocene (U–Th)/He ages (in the range of 10.0 ± 0.7–7.5 ± 0.4 Ma and 22.7 ± 1.3–19.1 ± 1.1 Ma, respectively) on maghemite/hematite from lateritic duricrust in Darling Range (Pidgeon et al., 2004) and on botryoidal goethite from metasandstone from Maynard Hills (Thern et al., 2011), both located in southern part of Western Australia. Moreover, the botryoidal goethite sample from Arkaroola in Southern Australia used here as an internal age standard, also yielded Miocene ages (20.8 ± 3.0 Ma). This suggests that the Miocene was the peak period, not only for CID formation in the Hamersley Province, but also for iron oxide precipitation and lateritisation in other parts of the Australian continent. This interpretation fits well with the geological and climatic record as well as with a number of previous age estimates based on other radiometric dating methods (see e.g., Idnurm, 1985; Bird and Chivas, 1989; Dammer, 1995; Pillans, 1998; Dammer et al., 1999; Anand and Paine, 2002; Vasconcelos et al., 2008 and references therein; 2013). 5. Conclusions New (U–Th)/He data of two goethite/hematite samples from different depths in a CID core in Robe River (Western Australia)

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constrain the timing of iron oxide formation and enable us to better understand the genesis of the CID. The most important results are summarised as follows: (i) (U–Th)/He ages range from 25.7 ± 0.6 to 2.2 ± 0.3 Ma and despite a high degree of scatter, the ages corroborate relationships expected from internal stratigraphy, where He ages of hematitic cores are older than ages of goethitic cortices and these are in turn older than ages of ferruginised wood in the matrix. The relative ages can be considered geologically meaningful and most likely represent the time of formation of iron oxides at Robe River in the Miocene, provided the (U–Th)/He system remained closed after mineral formation. (ii) It was possible to differentiate the following ore formation phases from the data: (i) formation of ooid cores filling the meandering palaeo-river channels latest during Early to Middle Miocene times between ~14.3 ± 3.7 Ma and 18.3 ± 3.5 Ma; (ii) Formation of goethitic cortex cementing the ooid hematitic cores in the late Middle to early Late Miocene at 11.6 ± 3.0 Ma; (iii) ferruginisation of the wood fragments in the matrix in the Late Miocene at ~10–7 Ma. (iii) Relatively high scatter of He ages and limited resolution of the (U–Th)/He dating method did not allow us to refine existing constraints on the climatic and weathering history of Western Australia. The only conclusion in this respect that can be made with confidence is that the unique environmental conditions favourable for CID formation existed in the Miocene. (iv) We confirm the observation of Vasconcelos et al. (2013) that the temperature applied for He extraction on iron oxides has a critical impact on the behaviour of parent nuclides. Traditionally applied laser heating at ~1000 °C likely induces loss of U and Th, which will result in erroneously old and more scattered ages. A modest extraction temperature (b 500 °C) generated by low-power lasers or preferably well-controllable resistance furnaces is recommended.

Acknowledgements We thank M. Verrall for help with the SEM, C. Scadding and A. Thomas for assistance with ICP-MS, A. Weisheit for providing the goethite sample from Arkaroola for internal calibration, and R. Morris for constructive suggestions on the CID genesis. P. Vasconcelos is thanked for sharing his knowledge on He extraction temperatures in Fe-oxides and for his constructive review. This project was funded by the Minerals Down Under Flagship, CSIRO. MD received financial support from the WA Geothermal Centre of Excellence. K. Metzger provided editorial handling. References Anand, R.R., Paine, M., 2002. Regolith geology of the Yilgarn Craton, Western Australia: implications for exploration. Australian Journal of Earth Sciences 49, 3–162. Bähr, R., 1987. Das (U + Th)/He-System von Hämatit als Chronometer für Mineralisationen. (Dr. rerum naturalium Thesis) Ruprecht-Karls University Heidelberg, Heidelberg. Bähr, R., Lippolt, H.J., Wernicke, S., 1994. Temperature-induced 4He degassing of specularite and botryoidal hematite: a 4He retentivity study. Journal of Geophysical Research 99, 17695–17707. Bersani, D., Lottici, P.P., Montenero, A., 1999. Micro-Raman investigation of iron oxide films and powders produced by sol–gel synthesis. Journal of Raman Spectroscopy 30, 355–360. Bird, M.I., Chivas, A.R., 1989. Stable isotope geochronology of the Australian regolith. Geochimica et Cosmochimica Acta 53, 3239–3256. Blackburn, T.J., Stockli, D.F., Walker, J.D., 2007. Magnetite (U-Th)/He dating and its application to the geochronology of intermediate to mafic volcanic rocks. Earth and Planetary Science Letters 259 (3–4), 360–371. Boschmann, W., 1986. Uran und Helium in Erzmineralien und die Frage ihrer Datierbarkeit. Heidelberger Geowissenschaftliche Abhandlungen, 4. University of Heidelberg, Heidelberg.

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