This article appeared in a journal published by Elsevier. The attached copy is furnished to the author for internal non-commercial research and education use, including for instruction at the authors institution and sharing with colleagues. Other uses, including reproduction and distribution, or selling or licensing copies, or posting to personal, institutional or third party websites are prohibited. In most cases authors are permitted to post their version of the article (e.g. in Word or Tex form) to their personal website or institutional repository. Authors requiring further information regarding Elsevier’s archiving and manuscript policies are encouraged to visit: http://www.elsevier.com/copyright
Transcript
This article appeared in a journal published by Elsevier. The attachedcopy is furnished to the author for internal non-commercial researchand education use, including for instruction at the authors institution
and sharing with colleagues.
Other uses, including reproduction and distribution, or selling orlicensing copies, or posting to personal, institutional or third party
websites are prohibited.
In most cases authors are permitted to post their version of thearticle (e.g. in Word or Tex form) to their personal website orinstitutional repository. Authors requiring further information
regarding Elsevier’s archiving and manuscript policies areencouraged to visit:
CO2 degassing and trapping during hydrothermal cycles relatedto Gondwana rifting in eastern Australia
I. Tonguc� Uysal a,⇑, Suzanne D. Golding b, Robert Bolhar b, Jian-xin Zhao c,Yue-xing Feng c, Kim A. Baublys d, Alan Greig e
a CO2CRC and Queensland Geothermal Energy Centre of Excellence, The University of Queensland, Queensland 4072, Australiab CO2CRC and School of Earth Sciences, The University of Queensland, Queensland 4072, Australia
c Radiogenic Isotope Facility, Centre for Microscopy and Microanalysis, The University of Queensland, Queensland 4072, Australiad School of Earth Sciences, The University of Queensland, Queensland 4072, Australia
e School of Earth Sciences, The University of Melbourne, Parkville, Victoria 3010, Australia
Received 2 February 2011; accepted in revised form 14 June 2011; available online 23 July 2011
Abstract
Intensive carbonate and clay mineral authigenesis took place throughout the Late Permian Bowen–Gunnedah–Sydneybasin system in eastern Australia. We conducted isotopic and trace element analyses of carbonate and clay minerals fromclastic sedimentary rocks of the Gunnedah Basin and the Denison Trough in the Bowen Basin. Rb–Sr isochron age dataof the illitic clays are consistent with episodic hydrothermal fluid flow events that occurred in association with Gondwanarifting accompanied by alkaline magmatism at �85 Ma and �95 Ma. Stable isotope data of carbonate and clay minerals fromthe Gunnedah Basin are indicative of meteoric waters from a high-latitude environment as the main fluid source, whereastrace element, Sr and Nd isotope data highlight mixing of meteoric fluids with magmatic and/or crustal components, witha possible input from marine carbonates for some samples. Trace metals, oxygen and strontium isotopes of dawsonites fromthe Denison Trough are interpreted to have been mobilised by fluids that interacted with evolved clastic sedimentary and mar-ine carbonate end members. According to the carbon isotope data, CO2 for calcite and ankerite precipitation was sourcedmainly from thermal degradation of organic matter and magmatism, whereas the CO2 used for dawsonite formation isinferred to have been derived from magmatic and marine sources. In the low permeability environments (particularly in coalseams), the increasing accumulation and oversaturation of CO2 particularly promote the precipitation of dawsonite.� 2011 Elsevier Ltd. All rights reserved.
1. INTRODUCTION
Continental hydrothermal systems typically occur inextensional tectonic settings that provide fracture pathwaysfor circulation of meteoric waters and CO2 from mantledegassing and dissolution of crustal carbonate deposits(Taylor, 1990; Mutlu et al., 2008; Uysal et al., 2009). Circu-lation of CO2-saturated, buoyant fluids gives rise to convec-tive heat transfer that controls the operation of continental
geothermal systems (Chiodini et al., 2007). CO2 accumu-lates partly in reservoir traps, with growing overpressurespotentially triggering seismic activity or hydrothermal erup-tions (Chiodini et al., 2004; Uysal et al., 2007a, 2009). Alter-natively, the pressure build-up may be limited by reactionof CO2 and water with Ca–Al-silicates to form carbonatesand phyllosilicate minerals (e.g., Huang and Longo, 1994).
Eastern Australian sedimentary basins provide an excel-lent natural laboratory for understanding the origin andevolution of CO2-bearing fossil hydrothermal fluids andthe mechanism of natural CO2 degassing and trapping. Sig-nificant CO2 production and mineral trapping occurred ineastern Australia during the Mesozoic, which is manifestedby widespread authigenic carbonate mineral occurrences
0016-7037/$ - see front matter � 2011 Elsevier Ltd. All rights reserved.
Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
throughout the Late Permian Bowen–Gunnedah–Sydney(BGS) basin system (Baker et al., 1995; Uysal et al.,2000a; Golab et al., 2006). In the Bowen Basin, carbonate– clay mineral authigenesis was a result of meteoric hydro-thermal processes controlled by episodes of Mesozoicextensional tectonism (Uysal et al., 2000a,b). Particularly,the occurrence of dawsonite provides clear evidence of thehigh CO2 partial pressure in sandstone reservoirs and thepresence of Na and Al released by dissolution of alumino-silicate minerals (e.g., Baker et al., 1995; Moore et al.,2005). Dawsonite has been observed in the Denison Troughof the Bowen Basin and the Gunnedah and Sydney Basins(Baker et al., 1995; Golab et al., 2006). Geochemical inves-tigation of this and other carbonate precipitates can be use-ful for assessing the evolution of CO2-rich subsurface fluids,which is important, among other things, for understandingthe fate of geo-sequestrated CO2 in deep sedimentary for-mations (Bachu, 2000).
The aim of the present study is to determine the originand evolution of CO2-rich fluids, and the mechanisms andtime scales of natural CO2 storage in some of the easternAustralian sedimentary basins. Constraints are derivedfrom clay mineralogy and isotopic and trace element anal-yses of authigenic carbonate and clay minerals from clasticsedimentary rocks of the Gunnedah Basin and the DenisonTrough (Bowen Basin). Geochemical and isotopic recordsgained from fossil CO2-rich natural systems, as we demon-strate in this study, can provide a test bed to assess the po-tential long-term effects of injection of anthropogenic CO2,which are otherwise difficult to evaluate in a laboratory-based simulation. Moreover, trace elements together withstable (O, H, C) and radiogenic (Sr, Nd) isotope composi-tions of hydrothermal mineral precipitates are excellentindicators of thermal and fluid flow histories, providingessential information on the CO2 degassing and trappingprocess in sedimentary reservoirs.
1.1. Regional geology
The BGS basin system extends north–south for ca.2000 km from central Queensland to southern New SouthWales in eastern Australia and comprises three intercon-nected sedimentary basins that are the Bowen, Gunnedahand Sydney Basins (Fig. 1A, B and D). The Bowen Basinand Gunnedah Basins are overlain in southern Queenslandand northern New South Wales by the Surat Basin (Figs.1C, 2), which is a component of the larger Mesozoic GreatArtesian Basin system.
The Gunnedah Basin represents the central part of theBGS basin system in northern New South Wales (Fig. 1Aand B) and was formed as a foreland basin containing Per-mo-Triassic marine (1200 m) and non-marine siliciclastic(4000 m) rocks (Tadros, 1993). These rocks are underlainby metamorphosed Paleozoic rocks of the Lachlan FoldBelt and abut the New England Fold Belt along the easternmargin of the BGS. The Gunnedah and Bowen Basin stra-tigraphy is more or less continuous across the Moree High;however, different names are applied to the same strati-graphic units in the Gunnedah and southern Bowen Basin(Fig. 2). The Gunnedah Basin contains coals seams that
are prospective for coal seam methane but have variablyhigh CO2 contents.
The Bowen Basin is a back-arc extensional to forelandbasin in the northern part of the BGS basin system andhas two major depocentres: the western Denison Troughand eastern Taroom Trough. The Bowen Basin is boundedto the east by the New England Fold Belt and underlain byPalaeozoic basement rocks. The Denison Trough consistsof north-trending half grabens (Fig. 1D) and contains upto 6500 m of Permo-Triassic marine and non-marine silici-clastic sedimentary rocks (Elliott, 1993). There are a num-ber of producing gas fields in the Denison Trough withhighest CO2 levels in the southern Denison fields up to33 mol%.
1.2. Sampling and analytical methodology
Clastic sedimentary rocks and carbonate veins weresampled from the Hoskisson and Bohena coal measuresnear the town of Narrabri in the Gunnedah Basin(Fig. 1). The sampling covered cored sections from the LatePermian Upper Black Jack Formation to the Early PermianMaules Creek Formation in boreholes Bibblewindi North1c (BWN-1C) and Bohena-12C (BOH) (Figs. 1C and 2;see Tadros, 1993 for a detailed lithostratigraphy). Somecore samples were also taken from Bohena South-2C(BS2C) (Fig. 1C). All representative lithology types (sand-stone, mudstone, claystone, pyroclastic rocks, fault gouge,and intrusions) were sampled at various depths with asmuch stratigraphic coverage as possible.
Petrographic analyses were carried out by X-ray diffrac-tion (XRD) and thin section studies of whole rock and veinsamples as well as by XRD study of clay mineral separates(<2 micron). The XRD analyses were carried out on an X-ray diffractometer with Bragg–Brentano geometry andCuKa radiation, operated at 40 kV and 30 mA at a scan-ning rate of 1�2h/min. The samples were prepared for sep-aration of the clay fraction by gently crushing the rocks tosand size, followed by disaggregation in distilled waterusing an ultrasonic bath. Different clay size fractions wereobtained by centrifugation, and the decanted suspensionswere placed on a glass slide. Following XRD analysis ofair-dried samples, the oriented clay-aggregate mounts wereplaced in an ethylene–glycol atmosphere at 30–40 �C over-night prior to additional XRD analyses. To determine illitecontent in illite–smectite mixed-layer clays, the method ofdifferential two-theta (D2h) was used, with an analytical er-ror of about ±5% (Moore and Reynolds, 1989).
ICP-MS trace element analyses were conducted for theclays and carbonate vein and cement samples. Dawsonitecements from the Denison Trough (samples provided byQueensland Museum), previously investigated by Bakeret al. (1995) for their oxygen and carbon isotopes, were ana-lysed in the current study by ICP-MS for their trace elementcontents. Carbonate samples were dissolved in 0.8 N nitricacid and centrifuged to remove impurities. Clay sampleswere dissolved with a mixture of HF and nitric acids on ahotplate, then evaporated to dryness, refluxed twice with ni-tric acid and dissolved in 2 N nitric acid. Aliquots of thesolutions were spiked with internal standards, diluted and
CO2 degassing and trapping in eastern Australia 5445
Author's personal copy
analysed on a Varian prototype 810 ICP-MS at the Univer-sity of Melbourne. The ICP-MS analytical protocol followsthat described by Eggins et al. (1997), except Tm was notused as an internal standard and W-2, a U.S.G.S. doleritestandard, was used for instrument calibration (Kamberet al., 2005). Other U.S.G.S. standards BHVO-1, BHVO-2 and BCR-2 were also run as unknowns to test accuracy.Our ICP-MS analytical protocol was based on the use ofmultiple internal standards for mass response corrections(Eggins et al., 1997), which ensures high-precision determi-nation of over 40 elements across the Periodic Table. In thisprotocol, 147Sm was used as internal standard to ensurehigh-precision measurements of REE contents and Sm–Nd ratio.
Illitic clay minerals were analysed for their stable isotope(d18O and dD) compositions. Oxygen was extracted from il-litic clays for isotope analyses using a CO2-laser and BrF5
(Sharp, 1990). Oxygen isotope values are reported in permil relative to V-SMOW and normalised to the interna-tional quartz standard (NBS-28) using a value of 9.6&.Replicate values for NBS-28 quartz (n = 6) analysed withthe samples had values that varied by less than 0.2&. Sam-ples and standards were heated overnight in a muffle fur-nace to 170 �C prior to loading into the vacuumextraction line to remove any adsorbed water. The sampleswere then evacuated for approximately 6 h and left over-night in a vapour of BrF5. Blank BrF5 was run until theyield was less than 0.1 micro moles oxygen. Oxygen waspassed through a fluorine-getter (in-line Hg diffusion pump)
and converted to CO2 by a graphite furnace; yields were re-corded and CO2 analysed on a Geo20–20 mass spectrome-ter at the GNS Laboratory, Lower Hutt, New Zealand.Hydrogen isotope analysis of clays was conducted at theUniversity of Queensland using a Thermo Delta V Advan-tage isotope ratio mass spectrometer coupled in continuousflow mode to a TC-EA with a zero blank solids autosam-pler. Hydrogen isotope compositions are reported in permil relative to V-SMOW, with an analytical uncertaintyof ±2& (1r). The Thermo Delta V Advantage isotope ratiomass spectrometer was calibrated and analytical precisionfor hydrogen isotope analysis established through replicateanalyses of NBS 30 (�65.7&) and IAEA CH7 (�100.3&).
Vein samples were handpicked and carbonate speciesand percentages were identified by XRD, which then deter-mined reaction temperature. Cement samples dominated bya single carbonate species based on XRD results were rununder conditions appropriate for that species. All sampleswere reacted off-line using the McCrea (1950) phosphoricacid digestion method. Calcite and dawsonite were reactedat 25 �C for 1 day, dolomite and ankerite at 50 �C for3 days and siderite at 75 �C for three days. Sample gaseswere analysed on an Isoprime dual inlet isotope ratio massspectrometer at the University of Queensland. Acid frac-tionation factors of 1.01025, 1.01066 and 1.00954 were usedfor the calculation of d18O values of the calcites, dolomitesand siderites, respectively (Sharma and Clayton, 1965;Rosenbaum and, Sheppard 1986). The calcite factor wasused for dawsonite in the absence of experimental data
Fig. 1. (A) Location of the BGS basin system in eastern Australia. (B) Simplified geological setting of Gunnedah Basin (C) Study area innorth Gunnedah Basin showing the major sub-basins and troughs and borehole locations. (D) Simplified geological settings of DenisonTrough. (E) Denison Trough showing borehole locations.
5446 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
for the dawsonite-acid fractionation factor (cf. Baker et al.,2005). The carbonate stable isotope analyses are reported inper mil (&) relative to V-SMOW for oxygen and V-PDBfor carbon, with analytical uncertainties better than±0.1& (1r). The Isoprime dual inlet isotope ratio massspectrometer was calibrated and analytical precision forcarbonates established through replicate analyses of NBS-18 and NBS-19.
Some selected carbonate samples from the GunnedahBasin were analysed for their 87Sr/86Sr ratios and a few of
them for their Sm/Nd isotopic systematics. Dawsonite ce-ments from the Denison Trough were also investigatedfor their 87Sr/86Sr ratios. Suitable illitic clay samples fromGunnedah Basin were dated by the Rb–Sr isochron tech-nique to determine the timing of major fluid flow events.
For radiogenic isotope analysis, calcite was dissolvedusing sub-boiling distilled 1 N acetic acid, and ankeriteand siderite were dissolved using distilled 1 N HCl at roomtemperature. Samples were then centrifuged and the leach-ates were dried and converted to chloride using sub-boiling
Fig. 2. Stratigraphy of the Gunnedah and Southern Bowen Basin (modified from Oil Company of Australia, 1985).
CO2 degassing and trapping in eastern Australia 5447
Author's personal copy
distilled HCl. For the Rb–Sr dating, illitic clay separateswere leached for 15 min at room temperature in 1 N dis-tilled HCl (Clauer et al. 1993). Leachate and residue wereseparated by centrifuging. The residue was rinsed repeat-edly with milli-Q water, dried and reweighed. Leachate, res-idue, and untreated samples were spiked with 85Rb–84Srmixed tracer and dissolved in a mixture of distilled HFand HNO3. The Sr-enriched fraction was separated usingcation exchange resins. Sr isotopic ratios were measuredon a VG Sector-54 thermal ionisation mass spectrometer(TIMS) in the Radiogenic Isotope Laboratory at Universityof Queensland. Sr was loaded in 1 N H3PO4 on a Ta singlefilament. Sr isotopic ratios were corrected for mass discrim-ination using 86Sr/88Sr = 0.1194. Long-term (6 years)reproducibility of the NBS SRM 987 (n = 442) is0.710249 ± 28 and the Pacific Ocean limestone standardEN-1 (n = 561) is 0.709176 ± 32 (2 r). During the courseof this study, the value obtained for the 87Sr/86Sr ratio ofstandard NBS 987 was 0.710229 ± 0.000008 (2r, n = 13).Rb–Sr isochron ages were calculated using the ISOPLOTprogram (Ludwig, 2003).
Dawsonite and calcite samples for Nd isotopic analyseswere selected, separated from the host rock and coal (daw-sonite cleats) and dissolved in dilute 1N HCl. The samplesolutions were then centrifuged to remove insoluble materialand other impurities and the clear solutions were then sepa-rated for Nd following standard cation exchange columnchemistry. Nd was separated from the REE fraction usingBio Bead� – HDEHP resin and eluting with 0.25 N HCl.Nd isotope ratios were corrected for mass fractionationusing 146Nd/144Nd = 0.7219. Nd isotopes were determinedusing the same TIMS as for Sr isotopes. Repeated measure-ments of La Jolla Nd standard on this mass spectrometeryield a mean 143Nd/144Nd ratio of 0.511860 ± 0.000018(2r). An in-house Nd metal standard yields a mean143Nd/144Nd ratio of 0.511969 ± 0.000025 (2r). The Ndprocedural blank is 25–50 pg.
2. RESULTS
2.1. Petrography
In thin section, the main detrital components of the sand-stones are quartz (mainly volcanic with some metamorphicgrains), volcanic rock fragments (commonly silicified or al-tered to kaolinite), K-feldspar (microcline, less sanidine),and some mica (usually altered to kaolinite or illite–smec-tite). Kaolinite, carbonate and silica are the main cementconstituents. Whole-rock powder XRD analyses indicatethat the carbonate minerals occurring as vein filling (cleatin coal) and cement in sandstones, mudrocks, intrusion andpyroclastic rocks in the Late Permian rocks are composedof siderite, calcite, ankerite, aragonite and dawsonite.
Two generations of siderite have been observed in theGunnedah Basin, which is similar to samples from the cen-tral/northern Bowen Basin (Uysal et al., 2000a). Siderite Irepresents a relatively early carbonate generation and isobserved mainly with kaolinite (Fig. 3A). It generally oc-curs in high abundance where it pervasively replaces almostthe entire rock content but is itself overprinted by kaolinite
(Fig. 3A). Siderite I can easily be recognised in thin sectionas it forms tiny crystal aggregates or radiating crystal clus-ters with relatively high relief. Siderite II is less commonand associated with calcite–ankerite cement. Calcite–anker-ite and siderite II have very similar optical and morphologicfeatures, so their distinction is based mainly on the XRDdata. They occur as clusters of well-formed small crystalsin thin section and are found mainly as a vein filling phaseboth at the macroscopic (several centimetres thick) andmicroscopic scale (Fig. 3B–D). Calcite–ankerite are also ob-served as replacement of volcanic rock fragments (Fig. 3E–G). The carbonate veining is locally very intense and com-monly associated with brecciation of large quartz grains(Fig. 3C and D). Dawsonite is observed filling thin (up to2 mm) veins in coal seams (Hoskisson Seam) in holesBWN 1C and BOH 12C. In hole BOH 2C it occurs togetherwith calcite and ankerite as a vein filling within a sandstone.
Illite–smectite commonly occurs as grain-coating, pore-lining and pore/fracture-filling cement and patchy replace-ment of volcanic rock fragments (e–h). In highly altered sam-ples (e.g., BOH 939.3 in Fig. 3), feldspars are not observeddue to complete dissolution followed by illite–smectite andcarbonate precipitation. We have developed a general para-genetic sequence of the authigenic minerals for Late Permianrocks in the Gunnedah Basin and Denison Trough based onour petrographical observations (Fig. 4).
2.2. Clay mineralogy
XRD analysis confirms that kaolinite is the dominantclay mineral in the <2-micron separates of all rock samples(Table 1). Kaolinite is the single clay phase or even the onlymineral in a number of rocks (claystone) between coalseams and in the immediate roof and floor sections of theHoskisson and Bohena Seams. It is most abundant in sam-ples from BWN-1C, whereas the samples from Bohena 12Cand Bohena 2C also contain significant mixed-layered illite–smectite (I–S).
Mainly two different types of I–S have been observedfrom the Gunnedah Basin samples. The first type of I–Sis characterized by a superstructure reflection at 26–28 Aand by diffraction peaks after ethylene glycolation near 9–9.5 and 13 A. These clays are short-range ordered(Reichweite, R = 1) I–S and contain 55–85% illite layers(see Moore and Reynolds, 1989). The second type of I–Sis R � 3 containing >90% illite layers. When glycolated,these clays show asymmetrical peaks at �9.8 A, with abroad shoulder near 11 A. A sample at 613.95 m from bore-hole BWN-1C is a smectite or randomly ordered (R = 0) I–S (17-A 001 peak with ethylene glycolation).
In sedimentary basins, the illite content in I–S usuallyincreases with increasing depth (Hower et al. 1976;Jennings and Thompson 1986). However, the composi-tion of I–S in BOH 12 is not systematically related todepth but rather shows irregular distribution, which isalso commonly observed in the Bowen Basin (Uysalet al. 2000a). For example, an illite-rich (>90% illite inI–S) R P 3 type I–S occurs at the top of the hole (UpperBlack Jack Formation), whereas the sample at the bot-tom is less illitic (75% illite in I–S) (Table 1).
5448 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
100 µm
calcite
100 µm
100 µm
Calcite-ankerite
siderite I
100 µm
100 µm
kaolinite
siderite I
A B
C D
100 µm
E
100 µm
F
100 µm
G H
calcite
calcite
I-S
I-S
I-S
I-S kaolinite
calcite
I-S
calcite
I-S
Fig. 3. Thin section micrographs showing clay and carbonate mineral authigenesis in the sandstones from the Gunnedah Basin: (A) Siderite I– kaolinite precipitation (sample BWN-1C at 647.2 m). (B) Calcite – ankerite micro-vein crosscutting a detrital quartz grain and earlierprecipitated siderite I (sample BWN-1C at 775.20 m). (C and D) Calcite-filled breccia veins (sample BOH-12C at 939.3 m). (E) Patchyreplacement of volcanic rock fragments by illite–smectite (I–S; white areas) and calcite. Note calcite postdates the I–S as it covers the areasthat were already replaced by I–S (sample BOH-12C at 939.3 m). (F) Grain-coating illite–smectite and replacement of a volcanic rockfragment by kaolinite and calcite. Note the kaolinite is covered by calcite indicating its earlier precipitation than the calcite (sample BOH-12Cat 939.3 m). (G) Illite–smectite coating a rock fragment that was replaced partly by calcite (sample BOH-12C at 939.3 m). (H) Illite–smectitefilling a micro-vein in a detrital quartz grain (sample BOH-12C at 939.3 m).
CO2 degassing and trapping in eastern Australia 5449
Author's personal copy
2.3. Oxygen, carbon and strontium isotope compositions of
carbonates
Oxygen and carbon isotopic compositions were deter-mined for twenty-six carbonate veins and cements fromthe Gunnedah Basin and nine dawsonite cements fromthe Denison Trough (Table 2). The Gunnedah Basin calciteand ankerite samples exhibit a broad range of d13C between�15.0 and 8.5 per mil and d18O between 5.8 and 25.1 permil that are quite similar to those of the calcites from thenorthern Bowen Basin (Uysal et al., 2000a) (Fig. 5A).The d18O values of the Gunnedah Basin siderites rangefrom 9.2 to 24.3 per mil. The d13C values vary from�14.5 to 3.1 per mil and are positively correlated with thed18O values. Although d18O and d13C values of sideritesfrom the Gunnedah and Bowen Basins show only limitedoverlap, they have parallel positive trends (Fig. 5A). TheGunnedah Basin dawsonites in the present study have a rel-atively narrow range of d13C between + 2.4 and �2.3 permil and d18O between 14.10 and 18.6 per mil. These d18Ovalues are generally higher than those of the calcite andankerite samples in the Gunnedah Basin. Dawsonite fromthe current study has similar but more restricted oxygenand carbon isotope compositions to the dawsonites re-ported earlier for some other locations in the GunnedahBasin as well as to those from the Denison Trough and Syd-ney Basin (Baker et al., 1995; Golab et al., 2006).
Strontium isotope results for 25 of the carbonate sam-ples are presented in Table 2. Three groups are recognisedfrom the Gunnedah Basin, which are characterised by dis-tinctively different ranges of 87Sr/86Sr ratios: 1) 0.70404 –0.70554; 2) 0.70664 – 0.70870; and 3) 0.72551 – 0.72557(Fig. 5B). Excluding the highly radiogenic samples fromgroup 3 and one sample with an extremely high d18O value(25.1 per mil), the remaining samples define a positive cor-relation between d18O values and 87Sr/86Sr ratios (Fig. 5B).Three dawsonite samples from the Gunnedah Basin have a
narrow range of 87Sr/86Sr ratios between 0.70857 and0.70870, falling into the second group of the Gunnedah Ba-sin samples. By contrast, 87Sr/86Sr isotope values of thedawsonites from the Denison Trough range between0.708569 and 0.731476, showing a clear negative correlationwith their d18O values (Fig. 5B).
2.4. Oxygen and hydrogen isotope compositions of illitic clay
minerals
Oxygen and hydrogen isotope analysis was carried outon nine illitic clay separates from four Gunnedah Basinsamples. The illitic clay minerals have a narrow range ofd18O values between 4.2& and 6.6& and dD values in arange from �86& to �120& (Table 3). We estimate forma-tion temperatures of 150 �C and 170 �C for the R = 1 andthe R P 3 type mixed layered I–S (Table 3), respectively(Pollastro 1993 and references therein). Using the oxygenand hydrogen isotope fractionation equations (Yeh, 1980;Sheppard and Gilg, 1996), we calculated oxygen and hydro-gen isotope compositions of fluids in equilibrium with theillitic clays ranging from �2.8& to �5.4& and from�65& to �101&, respectively (Fig. 6 and Table 3).
2.5. Trace element geochemistry
2.5.1. Carbonates
Trace element concentrations were determined for fif-teen carbonate samples from the Gunnedah Basin and ninedawsonite samples from the Denison Trough (Table 4).Post-Archean Australian Shale (PAAS)-normalised (Taylorand McLennan, 1985) rare earth element (REE) patternsfor the dawsonite and some calcite, ankerite and sideritesamples from the Gunnedah Basin and dawsonite samplesfrom the Denison Trough are presented in Fig. 7A–D.Dawsonites from the Gunnedah Basin differ significantlyin REE concentrations from samples from the Denison
Fig. 4. General paragenesis of common authigenic minerals in the Gunnedah Basin sediments.
5450 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
Trough, with the Gunnedah Basin dawsonites being highlyconcentrated in REE. One dawsonite sample from the Gun-nedah Basin is even more enriched in REE than PAAS(Fig. 7C). On the other hand one of the calcite samples withan unusually high 87Sr/86Sr ratio (BOH-12C-931.60) is sig-nificantly depleted of light REE (LREE) (Fig. 7A), imply-ing a different fluid source for this sample.
A well-defined positive correlation is evident betweenZr/Hf and Y/Ho ratios for carbonate vein samples and ce-ments from the Gunnedah Basin and dawsonite cementsfrom the Denison Trough (Fig. 8A). The majority of car-bonate samples from the Gunnedah Basin have Zr/Hf val-ues significantly higher than the chondritic value (>�35).Two dawsonite cleats from the Gunnedah Basin plot sepa-rately from the positive correlation trend in Fig. 8A due totheir extremely high superchondritic Zr/Hf ratios (see Table4), and these have been excluded from the regression line.Y/Ho correlates positively with d18O values of the DenisonTrough dawsonites and negatively with 87Sr/86Sr ratios ofthe Gunnedah Basin samples (Fig. 8B and C).
2.5.2. Illitic clay minerals
Trace element concentrations were determined for eightmixed layered illite–smectite clay separates prepared fromfour Gunnedah Basin samples (Table 5). These clay sepa-
rates have generally very high REE contents in comparisonto similar mixed layered illite–smectite clays from the Bo-wen Basin (see Uysal and Golding, 2003). Based onPAAS-normalised REE patterns and total REE contentsfrom the current study, illitic clay minerals from the Gunne-dah Basin can be grouped into 2 types: 1) samples withREE content higher than PAAS and no Eu negative anom-aly; and 2) those with lower REE content and also showinga pronounced negative Eu anomaly (Fig. 9; Table 5). A po-sitive trend can be seen between the extent of the Eu anom-aly ((Eu/Eu*)PAAS) and total REE contents of the illiticclays (Fig. 10A). There is an inverse relationship betweend18O and (Eu/Eu*)PAAS with samples clustering in 3 groups(Fig. 10B). There is also a negative correlation betweend18O values of the clays and total REE contents(Fig. 10C), ratios of light REE (LREE) to heavy REE(HREE), LREE to middle REE (MREE), and MREE toHREE (Fig. 10D–F). In addition, an inverse correlationis evident between d18O values and Y/Ho and Zr/Hf(excluding BOH 994.5 samples) ratios (Fig. 10G and H).
2.6. Rb–Sr geochronology of illitic clay minerals
Rb–Sr data for the untreated, leachates, and some acid-leached (residues) clay fractions from three sandstones and
Table 1Clay mineralogy data.
Sample-depth (m) Host rock Clay mineralogy Illite in I–S (%) I–S ordering
CO2 degassing and trapping in eastern Australia 5451
Author's personal copy
Tab
le2
Oxy
gen
,ca
rbo
nan
dst
ron
tiu
mis
oto
pe
dat
afo
rca
rbo
nat
esa
mp
les.
Sam
ple
-dep
th(m
)H
ost
rock
Min
eral
ogy
18O
(&V
-S
MO
W)
min
eral
13C
(&V
-PD
B)
min
eral
18O
(%,
V-S
MO
W)
wat
erat
80�C
18O
(%,
V-S
MO
W)
wat
erat
120
�C
18O
(%,
V-S
MO
W)
wat
erat
170
�C
87S
r/8
6S
r±
2r
Vei
n(G
un
ned
ahB
asin
)B
WN
-1C
-775
.20
mst
-sst
Cal
cite
,an
ker
ite
7.0
�15
.0�
12.4
�8.
1�
4.2
0.70
554
±08
BW
N-1
C-7
87.1
0m
st-s
stB
arit
e,ca
lcit
e8.
8�
0.6
�10
.6�
6.2
�2.
4B
WN
-1C
-837
.51
mst
Cal
cite
,m
ino
ran
ker
ite,
qu
artz
9.2
�6.
5�
10.2
�5.
9�
2.1
0.70
4275
±10
BW
N-1
C-8
44in
tC
alci
te,
trac
ean
ker
ite
6.6
�3.
3�
12.8
�8.
5�
4.7
0.70
4663
±08
BW
N-1
C-8
52.5
0in
tC
alci
te,
min
or
ank
erit
e10
.7�
1.5
�8.
7�
4.4
�0.
60.
7040
44±
08B
OH
-12C
-557
.30
mst
Cal
cite
,Q
uar
tz11
.0�
3.4
�8.
4�
4.1
�0.
30.
7067
56±
10B
OH
-12C
-698
.30
coal
Daw
son
ite
(cle
atin
coal
)18
.10.
2�
1.3
3.0
6.8
0.70
8569
±10
BO
H-1
2C-6
98.5
0co
alD
awso
nit
e(c
leat
inco
al)
17.3
�0.
2�
2.1
2.2
6.1
0.70
8618
±08
BO
H-1
2C-7
25.6
sst-
silt
stC
alci
te,
Qu
artz
,k
aoli
nit
e25
.14.
45.
610
.013
.80.
7067
37±
07B
OH
-12C
-931
.60
mst
Cal
cite
9.8
8.5
�9.
6�
5.3
�1.
50.
7255
12±
10B
OH
-12C
-935
.20
mst
Cal
cite
9.7
6.4
�9.
7�
5.4
�1.
60.
7255
68±
09B
OH
-12C
-991
.10
pyr
Cal
cite
,m
ino
rq
uar
tz10
.0�
2.4
�9.
5�
5.1
�1.
30.
7046
79±
10B
OH
-12C
-998
.5p
yrC
alci
te10
.67.
0�
8.8
�4.
5�
0.7
0.70
5452
±08
BS
-2C
-615
.40
coal
Ara
gon
ite
12.6
1.2
�6.
8�
2.5
1.4
0.70
7327
±10
BS
-2C
-632
sst
Cal
cite
,m
ino
rk
aoli
nit
e5.
8�
7.8
�13
.6�
9.3
�5.
40.
7066
39±
10B
S-2
C-7
18.1
0ss
tD
awso
nit
e,q
uar
tz,
calc
ite,
ank
erit
e14
.3�
1.7
�5.
2�
0.8
3.0
0.70
8702
±10
BS
-2C
-738
.5ss
tA
nk
erit
e,m
ino
rq
uar
tz13
.7�
10.8
�8.
9�
3.9
0.5
Cem
ent
(Gu
nn
edah
Bas
in)
BW
N-1
C-6
13.9
5m
stA
nk
erit
e10
.6�
14.5
�12
.0�
7.0
�2.
6B
WN
-1C
-643
.25
mst
Sid
erit
e15
.20.
4�
6.4
�1.
62.
80.
7073
78±
08B
WN
-1C
-647
.2ss
tS
ider
ite
13.5
�7.
5�
8.1
�3.
31.
1B
WN
-1C
-775
.20
mst
-sst
Sid
erit
e24
.33.
12.
77.
511
.8B
OH
-12C
-557
.30
mst
Sid
erit
e18
.0�
0.7
�3.
61.
25.
5B
OH
-12C
-604
sst
Sid
erit
e17
.6�
3.5
�4.
00.
95.
2B
OH
-12C
-704
.5ss
tA
nk
erit
e9.
8�
8.5
�12
.8�
7.8
�3.
4B
OH
-12C
-939
.30
sst
Cal
cite
8.2
�4.
1�
11.2
�6.
9�
3.1
BS
-2C
-632
sst
Sid
erit
e9.
2�
14.5
�12
.3�
7.5
�3.
2
Cem
ent
(Den
iso
nT
rou
gh)
ED
1-7
sst
Daw
son
ite
16.5
0.2
�2.
91.
45.
20.
7135
97±
10M
Y2-
2ss
tD
awso
nit
e14
.92.
3�
4.5
�0.
23.
60.
7084
54±
12M
Y2-
6ss
tD
awso
nit
e15
.11.
8�
4.3
0.0
3.8
0.70
7885
±08
SP
1-2
sst
Daw
son
ite
15.8
1.9
�3.
60.
74.
50.
7101
22±
08Y
B2-
11ss
tD
awso
nit
e13
.21.
0�
6.2
�1.
91.
90.
7126
30±
10Y
B3-
6ss
tD
awso
nit
e9.
90.
9�
9.5
�5.
2�
1.4
0.73
1476
±42
YB
3-11
sst
Daw
son
ite
10.1
1.1
�9.
3�
5.0
�1.
20.
7157
00±
10Y
B4-
3ss
tD
awso
nit
e12
.00.
9�
7.4
�3.
10.
70.
7181
72±
12Y
B4-
8ss
tD
awso
nit
e11
.90.
5�
7.5
�3.
20.
60.
7234
08±
12
sst
=sa
nd
sto
ne,
ms
t=
mu
dst
on
e,si
lst
=si
ltst
on
e,in
t=
intr
usi
on
,p
yr=
pyr
ocl
asti
c.F
luid
18O
com
po
siti
on
sw
ere
calc
ula
ted
usi
ng
the
calc
ite–
wat
er(O
’Nei
let
al.,
1969
;al
sofo
rd
awso
nit
e,se
eB
aker
etal
.,19
95),
do
lom
ite–
wat
er(f
or
ank
erit
e;O
’Nei
let
al.,
1969
;S
hep
par
dan
dS
chw
arcz
,19
70),
and
sid
erit
e–w
ater
(Car
oth
ers
etal
.,19
88)
frac
tio
nat
ion
s.
5452 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
one fault gouge collected from different boreholes andstratigraphic levels in the Gunnedah Basins are presentedin Table 6 and on Fig. 11. The data show well-defined linearrelationships corresponding to an isochron age of93.90 ± 0.95 Ma (initial 87Sr/86Sr = 0.70566 ± 0.00014) forsamples BOH 994.5 and BWN 831 (excluding 1–0.2 lmand <2 lm fractions with trace K-feldspar content) (Table6; Fig. 11A). The data of untreated clays and the leachatesonly (excluding the residue) define a linear relationship withan analytically indistinguishable age of 94.4 ± 5.6 Ma (ini-tial 87Sr/86Sr = 0.70564 ± 0.00036) (Fig. 11B). The Rb–Srisotope data for the <0.2 lm and 2–0.2 lm fractions ofsample BOH 939.3 show a linear relationship correspond-ing to an age of 85.7 ± 3.0 Ma (initial 87Sr/86Sr =0.7276 ± 0.0017) (Fig. 11C). An individual leachate, resi-due, and untreated line for the <0.2 lm fraction of sampleBS2C 632 (Table 6) gives an apparent age with a very largeerror (84 ± 13 Ma), with an initial 87Sr/86Sr value of0.7091 ± 0.0038 (Fig. 11D).
2.6.1. Sm/Nd isotope data for carbonates
Sm–Nd isotope analysis was carried out on five carbon-ate samples from the Gunnedah Basin, and the data are re-ported in Table 7 and on Fig. 12. In the case of samplesBOH698.30, BOH698.50 and BS718.10, three aliquots weremeasured resulting in a total of nine analyses, while sampleBOH935.20 and BOH931.60 were analysed twice and once,respectively (Table 7). Analyses for subsamples were com-bined to calculate average isotopic compositions for eachsample, and these data are shown on Fig. 12.
Sm and Nd concentrations vary by more than two or-ders of magnitude. 147Sm/144Nd ratios are variable, rangingfrom 0.192 to 0.797, with sample BOH931.60 yielding ahighly anomalous 147Sm/144Nd ratio of 0.797. Using aninferred crystallisation age of 85 Ma obtained from theRb–Sr isochrons of associated clay minerals (referFig. 11C) and measured 143Nd/144Nd values, we calculatedthe Nd isotopic composition at 85 Ma. The initial Nd isoto-pic compositions are expressed in the epsilon (e) notation,calculated after McCulloch and Wasserburg (1978) using147Sm/144Nd CHUR, present = 0.1967 (Jacobsen and Was-serburg, 1980) and 143Nd/144Nd CHUR = 0.512638 (Gold-stein et al., 1984). eNd (85 Ma) values are highly variable,ranging from �1.9 to 45.5 units (Table 7, Fig. 12A).
3. DISCUSSION
3.1. Clay mineralogy
The abundance of kaolinite throughout entire strati-graphic sections from Early to Late Permian in severalGunnedah Basin holes indicates that acidic fluids may haveaffected the whole system. Two main processes, abundantCO2 degassing with the associated carbonic acid generationand intense meteoric water flow flushing, may supply diluteand acidic solutions for kaolinite precipitation (cf., Batta-glia et al., 2007). As indicated by the stable isotope system-atics (see below), both mechanisms were responsible for theformation of the clay – carbonate mineral association in theinvestigated core samples. CO2-bearing fluids reacting with
feldspars, volcanic rock fragments, mica and some Fe–Mgsilicates will release an excess of reaction products such asCa2+, Na+, K+, Mg2+, Fe2+ and Al3+. At the earliest stageof the reaction, kaolinite will precipitate from HCO3
�1-richacidic fluids. As the reaction proceeds, depending on theavailability of cations (particularly K+), precipitation of il-lite–smectite will follow in a slightly acidic to neutral envi-ronment (e.g., Barclay and Worden, 2000; Battaglia et al.,2007). The remaining, more cation-rich fluids will lead todeposition of carbonate minerals and silica (e.g., Barclayand Worden, 2000; Kaszuba et al., 2006) (see below).
In the investigated Gunnedah Basin core samples, illite–smectite occurs as randomly ordered (R = 0-type, <10% il-lite), regularly ordered (R = 1-type, 60–85% illite) and long-range ordered (R P 3-type, >90% illite) mixed-layered clayminerals. The formation of such illitic clay minerals is con-trolled mainly by temperature, although there are also someother factors affecting the illite formation, such as chemicalcomposition of fluids and water–rock ratios related to per-meability (e.g., Uysal et al., 2000c). By analogy with litera-ture data (Pollastro, 1993 and references therein), theoccurrence of different types of illite–smectite in the studyarea indicate a temperature range of formation fromroughly 80 �C (smectite-rich) to 170 �C (illite-rich). How-ever, the distribution of different illitic clay types in theinvestigated holes is not depth-related and changes irregu-larly within relatively brief stratigraphic intervals. Similarvertical variability and zonation of clay minerals observedfor illitic clays from the Bowen Basin was attributed tochanges in fluid physico-chemical characteristics controlledlargely by variable fluid/rock ratios in a hydrothermal sys-tem (Uysal et al., 2000b,c). In many ways, composition andthe distribution patterns of clay minerals from the Gunne-dah Basin are similar to those of the clays in the Bowen Ba-sin, which is not unexpected as both basins have beenaffected by regional-scale extensional tectonism and theassociated hydrothermal processes.
3.2. Implications of Rb–Sr geochronology of illitic clay
minerals
Clay size separates of Late Permian clastic rocks in theGunnedah Basin display well-developed linear data arrayson a Rb–Sr isochron diagram. Such linear relations canrepresent either a result of mixing between an older detritalillite population and a younger authigenic illite–smectitegeneration, or complete isotopic equilibration of the entireauthigenic clay population. In the former case, the linearrelationship between 87Rb/86Sr and 87Sr/86Sr would repre-sent a mixing line of two end members with no meaningfulage information, whereas the latter relation would providean isochron whose slope yields the age of illitic clay gener-ation. X-ray diffraction patterns of clay size fractions showthat clays of all different size fractions consist of the sametype of mixed-layer illite–smectite with no discrete illitebeing observed as a potential detrital contaminant. Basedon the X-ray diffraction patterns, whole rock powder sam-ples from which clays for Rb–Sr dating were separated arefree from feldspars, which were dissolved and alteredcompletely to clay and carbonate minerals. Thin section
CO2 degassing and trapping in eastern Australia 5453
Author's personal copy
analyses show that the samples contain newly grown authi-genic illite–smectite evidenced by their common occurrenceas pore- and fracture-filling mineral and grain-coatings(Fig. 3F–H) (cf., Boles, 2008). Illite–smectite may also haveformed partly as replacement of primary mica through acombination of transformation and dissolution–precipita-tion and hence the clay fractions may contain some inher-ited 2:1 layers. However, the linear relationships inFig. 11 suggest that any primary isotopic signatures havebeen reset and the Rb–Sr data most likely represent iso-chron ages of the clay precipitation.
Rb–Sr age data of the illitic clays coincide with the tim-ing of Gondwana rifting episodes in eastern Australia. Theextension between Australia and Antarctica and the open-ing of Tasman Sea occurred as episodic events, associated
with alkaline magmatism in New Zealand and southeastAustralia (Veevers et al., 1991; Waight et al., 1998). Onthe western margin of the rift, widespread uplift and ero-sion of the Australian Eastern highlands at �96 Ma repre-sent the stages of crustal breakup and extension (O’Sullivanet al., 1995; Veevers et al., 1991), which is consistent withthe timing of illite formation in samples BOH 994.5 andBWN 831.2 (Fig. 11A and B). Intraplate alkaline igneouscentres recognised in both Australia and New Zealand arepossibly linked with the latest Albian to early Cenomanian(100–95 Ma) rifting event (O’Sullivan et al., 1995 and refer-ences therein). A later episode of rifting is demonstrated bythe first appearance of oceanic crust in the Tasman Sea at�84 Ma (Veevers et al., 1991), with the emplacement ofwithin-plate alkaline magmatism (Waight et al., 1998).The Rb–Sr age of 86 ± 3 Ma for the illite generation insample BOH 939.3 (Fig. 11C) is in agreement with the tim-ing of the later tectonism and magmatism.
3.3. Origin and evolution of fluids
3.3.1. Stable isotope results
As discussed above, interaction between CO2-rich richhydrothermal fluids and siliciclastic rocks will enhance thecation release leading to formation of a widespread clay,carbonate and quartz mineral association (Huang andLongo, 1994; Barclay and Worden, 2000; Kaszuba et al.,
Fig. 5. O–C–Sr isotope variation diagrams. (A) d13C vs. d18O forstudied calcite, ankerite, siderite and dawsonite samples and fordawsonites compiled from the literature. (B) 87Sr/86Sr vs. d18O forGunnedah Basin carbonates and Denison Trough dawsonites.
Table 3Oxygen and hydrogen isotope data for illitic clays from the Gunnedah Basin.
Fig. 6. Calculated d18O and dD values of fluids in equilibrium withillitic clay minerals in the Gunnedah Basin in comparison to thed18O and dD values of formation waters reported for low and high-latitude sedimentary basins (Clayton et al., 1966).
5454 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
Tab
le4
Tra
ceel
emen
tco
nce
ntr
atio
ns
for
carb
on
ate
fro
mth
eG
un
ned
ahB
asin
and
Den
iso
nT
rou
gh(p
pb
).
Sam
ple
Min
eral
Sc
Cr
Rb
Sr
YZ
rN
bC
sB
aL
aC
eP
rN
dS
mE
u
Gu
nn
eda
hB
asi
n
BO
H-5
57.3
Ca
12,6
8342
9916
3552
,652
46,5
5513
,100
8.4
276
25,7
2720
2084
6517
7211
,011
3717
923
BO
H-6
98.3
Dw
2827
27,2
9012
480
,836
145,
819
80,5
9230
5.5
577,
713
45,6
9814
6,87
119
,563
75,6
1325
,953
5944
BO
H-6
98.5
Dw
1141
21,9
4119
022
7,72
829
,541
51,3
8324
1397
9,64
986
2926
,047
3616
14,7
6249
2310
81B
OH
-725
.6C
a84
3815
1154
912
04,8
2710
,499
1214
1.7
8567
,011
8298
17,1
1321
0284
2120
6348
1B
OH
-931
.6C
a86
8826
518
95,7
7212
,285
245
0.7
0.2
75,1
9252
301
8691
314
6852
4B
OH
-939
.3C
a-ce
men
t52
3339
826
8060
,819
15,3
0917
4612
148
71,7
8813
,864
27,2
9744
6118
,888
5157
910
BO
H-9
98.5
Ca
5992
2068
895
87,9
3524
,986
1146
2218
789
5081
,926
346,
983
36,6
4416
8,40
639
,152
6402
BW
N-7
75.2
Ca-
An
k18
459
5.2
2343
,698
4899
2828
0.1
0.3
161,
679
3485
6675
1059
4618
1119
241
BW
N-8
44C
a50
8979
6.7
354,
429
4961
266
0.1
2.2
15,7
0462
8015
,919
2043
8441
1571
544
BW
N-8
52.5
Ca-
An
k53
611
611
773,
019
724
1751
0.1
1.2
24,0
0720
338
851
218
4418
BW
N-6
43.2
5S
id-c
emen
t12
7615
635
393
2039
9618
7015
186
16,0
3639
3079
5194
336
1679
315
4B
WN
-647
.2S
id-c
emen
t30
3820
511
7124
,322
7262
3273
2.7
151
43,0
3142
,367
96,5
0011
,411
40,4
7471
1210
84B
S-6
32S
id30
0315
3477
2151
,136
11,1
7111
313.
358
311
9,24
033
,124
82,1
6911
,044
44,4
9510
,508
418
BS
-718
.1D
w-C
a-A
nk
15,4
8911
8,41
848
966
9,77
437
,540
10,5
628.
835
338,
408
3179
9971
1811
10,9
4963
2327
49B
S-7
38.5
An
k15
,328
7046
1525
1,61
165
,073
187,
835
6.3
0.7
893,
278
7118
18,3
0524
4811
,580
4876
1675
Den
iso
nT
rou
gh
ED
1-7
Dw
-cem
ent
160
1831
922
2034
380
309
2.3
8714
0,98
691
199
2510
135
18M
Y2-
2D
w-c
emen
t88
1956
802
2453
271
513
5.5
9969
6046
111
1813
551
310
625
MY
2-6
Dw
-cem
ent
283
3127
924
4755
468
422
3.3
115
83,3
6226
263
482
334
8326
SP
1-2
Dw
-cem
ent
286
474
1003
1600
489
346
2.9
117
30,5
7251
411
2814
859
716
062
YB
2-11
Dw
-cem
ent
8252
211
5011
9019
157
17.
910
880
7913
027
333
120
265
YB
3-11
Dw
-cem
ent
284
494
1335
1315
272
481
1.9
128
1810
5312
116
6118
7Y
B4-
3D
w-c
emen
t25
439
116
2389
819
556
52.
215
932
4494
199
2488
226
YB
4-8
Dw
-cem
ent
457
525
1535
2219
285
641
3.4
151
2695
149
330
3813
934
10S
S5-
20D
w-c
emen
t58
014
0728
9463
2013
5999
115
265
13,8
5327
9163
8483
432
1866
813
5
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
UZ
r/H
fY
/Ho
Gu
nn
eda
hB
asi
n
BO
H-5
57.3
Ca
5147
908
6120
1417
4139
605
3695
535
951.
634
2817
7443
4213
8.5
32.8
5B
OH
-698
.3D
w34
,335
5649
31,5
6459
4714
,927
1871
9702
1248
103
8.4
2358
26,8
3220
,264
783.
124
.52
BO
H-6
98.5
Dw
6582
1040
5798
1116
2742
346
1818
247
826.
313
8784
2463
3462
8.7
26.4
6B
OH
-725
.6C
a20
7535
720
5339
610
0413
781
111
120
0.3
599
1022
254
60.6
26.5
1B
OH
-931
.6C
a49
3765
427
6844
490
599
509
692
0.4
503
193
1712
4.0
27.6
9B
OH
-939
.3C
a-ce
men
t53
2671
833
5555
812
4215
385
211
269
0.7
17,7
6316
0228
425
.427
.41
BO
H-9
98.5
Ca
25,5
8422
7677
0610
5420
5924
012
9617
027
2.4
951
242
942
.223
.70
BW
N-7
75.2
Ca-
An
k12
1314
970
512
326
629
145
1922
0.1
8687
6212
5.8
39.9
5B
WN
-844
Ca
1467
192
960
174
414
5432
048
30.
363
735
2794
.628
.55
BW
N-8
52.5
Ca-
An
k51
858
1548
850
96
0.0
480
3.7
1427
4.3
48.2
1B
WN
-643
.25
Sid
-cem
ent
773
133
804
162
460
7449
070
600.
616
1793
118
531
.324
.68
BW
N-6
47.2
Sid
-cem
ent
4568
541
2270
343
740
9553
669
118
0.4
12,5
9722
,120
819
27.8
21.1
7B
S-6
32S
id84
8510
5742
7658
010
7511
456
765
670.
953
,100
28,3
8635
5116
.919
.27
BS
-718
.1D
w-C
a-A
nk
9747
1621
8935
1656
3906
496
2751
381
991.
234
359
1771
910
6.6
22.6
7B
S-7
38.5
An
k11
,633
2307
12,9
6622
6849
0555
628
6938
215
041.
676
139
2929
8612
4.9
28.6
9
CO2 degassing and trapping in eastern Australia 5455
Author's personal copy
2006; Battaglia et al., 2007). Assuming the illitic clays andcarbonates were formed in a similar temperature rangefrom roughly 80 �C (smectite-rich) to 170 �C (illite-rich)and from fluids with a similar origin we can estimate theoxygen isotope composition of fluids from which the car-bonates precipitated (Table 2). We used model tempera-tures of 80 �C, 120 �C and 170 �C, and relevant mineral–water fractionations to calculate the oxygen isotope compo-D
enis
on
Tro
ug
h
ED
1-7
Dw
-cem
ent
7615
8515
364.
727
3.8
130.
234
1814
534
24.3
24.8
MY
2-2
Dw
-cem
ent
8813
6311
263.
419
2.7
160.
253
8027
046
31.7
24.6
MY
2-6
Dw
-cem
ent
108
1910
720
465.
933
4.6
140.
229
4224
050
29.8
23.8
SP
1-2
Dw
-cem
ent
165
2612
722
496.
335
4.9
170.
234
4846
446
20.4
22.6
YB
2-11
Dw
-cem
ent
275
398
182.
212
1.7
200.
250
4688
2928
.524
.0Y
B3-
11D
w-c
emen
t38
1170
1227
3.3
182.
619
0.1
5919
122
3124
.722
.1Y
B4-
3D
w-c
emen
t27
642
820
2.5
152.
123
0.1
4465
4323
24.4
23.5
YB
4-8
Dw
-cem
ent
5212
6912
293.
621
2.9
230.
147
9610
028
28.1
23.0
SS
5-20
Dw
-cem
ent
557
7736
962
137
17.1
9613
.340
0.7
4908
1816
137
24.6
22.0
Ca
=ca
lcit
e,D
w=
daw
son
ite,
An
k=
ank
erit
e,S
id=
sid
erit
e.
Fig. 7. PAAS-normalised REE pattern for (A) calcite from theGunnedah Basin, (B) dawsonite from the Denison Trough, (C)dawsonite from the Gunnedah Basin, and (D) siderite and ankeritefrom the Gunnedah Basin.
5456 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
sition of fluids in equilibrium with calcite, ankerite, sideriteand dawsonite from the Gunnedah Basin and dawsonitefrom the Denison Trough (Table 2). A low temperature for-mation (80 �C or lower) of calcite and ankerite (one sider-ite) samples is unlikely because this would implyunrealistically low averaged fluid oxygen isotope composi-tions of 6�11& (cf. Uysal et al., 2000a). Considering ahigher temperature range between 120 and 170 �C wouldgive more realistic averaged oxygen isotope compositionsin a range between �6& and �2& for the calcite andankerite samples. These oxygen isotope compositions are
similar to those calculated for the fluids of the illitic clays(see below). Dawsonite formation requires lower tempera-tures (<120 �C) if similar meteoric-dominated waters wereinvolved in the precipitation of this mineral (see also Bakeret al., 1995). We interpret the siderite I as an early diage-netic mineral precipitated as a result of fermentation reac-tions of organic matter at temperatures less than 80 �C,which is also common in the Bowen Basin (Uysal et al.,2000a).
The calculated oxygen and hydrogen isotope composi-tions of fluids in equilibrium with the illites are significantlylow (�2.8& to �5.4& and from �65& to �101&, respec-tively), compared to the oxygen and hydrogen isotope com-positions of formation waters reported for most mid – tohigh-latitude sedimentary basins (Fig. 6; Clayton et al.,1966) and to those attained during burial diagenesis in othersedimentary basins (Clauer and Chaudhuri, 1995 and refer-ences therein). Oxygen and hydrogen isotope compositionsof waters from which the illitic clays in the Gunnedah Basinprecipitated are comparable to geothermal waters reportedfor most high-latitude environments (Truesdell andHulston, 1980). In this context geothermal waters are gener-ally enriched in 18O relative to meteoric water due to isoto-pic exchange between hot waters and rocks or mixing withmagmatic waters (Truesdell and Hulston, 1980) (Fig. 6).By contrast, the dD compositions of meteoric hydrothermalwaters are not changed significantly unless they are affectedby phase changes (Truesdell and Hulston, 1980) or extremefluid/rock interactions (Nesbitt et al., 1989). This dissimilar-ity between the oxygen and hydrogen isotope systematics isdue to the different amounts of oxygen and hydrogen in therock and in the fluid, with the fluid hydrogen content beingoverwhelmingly dominant in most systems (Sheppard,1986). Therefore, hydrogen isotope data provide the best re-cord of the initial isotopic composition of fluids, providedthere is no post-formation hydrogen isotope exchange withfluids related to a younger fluid flow event (e.g., Longstaffeand Ayalon, 1990). We consider a post-formation hydrogenisotope exchange process unlikely because this would alsohave affected the linear data arrays on Rb–Sr isochron dia-grams of the illitic clays, which is, however, not the case (seeFig. 11). We estimated the initial oxygen isotope composi-tion of meteoric water from the hydrogen isotope composi-tion of the fluid and the meteoric water line (MWL, Fig. 6)assuming the present day MWL equation. The calculatedhydrogen isotope composition of fluids in equilibrium withthe illitic clays in the Gunnedah Basin range from �65&
to �101& (Table 3). These values would give an initial oxy-gen isotope composition for the fluid of about �9& to�15&, which is consistent with the relatively high latitudeposition of east-southeast Australia during clay formationin the Cretaceous (Veevers, 1984). However, dD values inTable 3 are somewhat variable that may indicate local mix-ing with magmatic waters (Truesdell and Hulston, 1980), oralternatively erroneously estimated paleotemperatures usedfor calculation of the dD values. Indeed, trace element,Sr and Nd isotope data indicate that both carbonate andillite-precipitating fluids also incorporated magmatic as wellas crustal and possibly marine components, as discussedbelow.
Fig. 8. Plots of Y/Ho vs Zr/Hf, d18O, and 87Sr/86Sr for GunnedahBasin calcite, ankerite, siderite and dawsonite and Denison Troughdawsonite. (A) A positive correlation between Y/Ho and Zr/Hf isevident for both Gunnedah Basin (two dawsonite samples withanomalous Zr/Hr ratios were excluded) and Denison Troughsamples. (B) Denison Trough dawsonite samples show somecorrelation between their Y/Ho and d18O values. (C) Somecorrelation is apparent between Y/Ho and 87Sr/86Sr for GunnedahBasin carbonates.
CO2 degassing and trapping in eastern Australia 5457
Author's personal copy
3.3.2. Trace element data: carbonates
Superchondritic Zr/Hf ratios for the Gunnedah Basincarbonates imply that Zr was mobile in the carbonate pre-cipitating fluid, although Zr is generally considered to beone of the most immobile elements. High-field strengthincompatible elements such as Zr, Y and REE are enrichedin alkaline igneous rocks and further concentrated andtransported in associated orthomagmatic hydrothermal flu-ids through fluoride complexes (e.g., Salvi and Williams-jones, 1990; Wood, 1990; Salvi and WilliamsJones, 1996).Superchondritic Zr/Hf ratios of the Gunnedah Basin car-bonates with their relatively high REE contents can be ex-plained by mixing of deeply circulating meteoric waterswith alkaline magmatic fluids and transport of the mixedfluids to the coal measures where the carbonates were pre-cipitated. The positive correlation between Zr/Hf and Y/
Table 5Trace element concentration of illitic clay fractions from the Gunndedah Basin.
BOH-939.3 994.5 BOH-994.5 BWN-1C-831.2 BWN-1C-831 BWN-1C-831.2 BS-632 BS-6322–0.2 m <0.2 m 0.5–0.2 m <0.2 m <2 m 1–0.2 m <2 m <0.2 m
Fig. 9. PAAS-normalised REE pattern for illitic clay mineralsfrom the Gunnedah Basin.
5458 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
Fig. 10. Illitic clay minerals from the Gunnedah Basin display a positive and negative correlation between Eu anomaly and total REE content(A) and d18O (B), respectively. Negative correlations are shown between d18O and total REE content (C), LREE/HREE (D), LREE/MREE(E), and MREE/HREE (F) ratios. d18O values of the clays display also inverse correlations with their Y/Ho (G) and Zr/Hf (excluding BOH994.5 samples) (H) ratios.
CO2 degassing and trapping in eastern Australia 5459
Author's personal copy
Ho ratios for the carbonates further support the transportof Zr and Y as fluoride complexes in fluids from whichthe dawsonite, calcite, ankerite and some siderite precipi-tated in the Gunnedah Basin. The dawsonite in the DenisonTrough, however, show no superchondritic Zr/Hf ratios(Table 4), which indicates that fluid sources were somewhatdifferent from those in the Gunnedah Basin.
The negative correlation between 87Sr/86Sr and Y/Hofor the Gunnedah Basin and positive correlation betweend18O values and Y/Ho for the Denison Trough (Fig. 8Band C) samples provide more elaborate constraints on car-bonate fluid sources. Y/Ho fractionation is common in sea-water and hydrothermal fluids, reflected by elevated Y/Horatios in marine carbonates and hydrothermal deposits(Bau, 1996). In hydrothermal fluids, Y/Ho fractionation oc-curs through complexation with fluoride, which is higherfor Y than Ho resulting in increasing Y/Ho during migra-tion of fluorine-bearing hydrothermal fluids (Bau and Dul-ski, 1995). Increasing Y/Ho ratios of the Gunnedah Basincarbonates are associated with less radiogenic (or lower)87Sr/86Sr values, which is consistent with the above inter-pretation proposing mixing of meteoric waters with alkalinemagmatic fluids and transport of the mixed fluids to theshallower carbonate precipitation sites. We interpret thehigher Y/Ho ratios with increasing d18O values for the Den-ison Trough dawsonites to reflect increasing input frommarine sources. Our interpretations above are further sup-ported by the correlation between d18O and 87Sr/86Sr val-ues, which is positive for the Gunnedah Basin carbonatesbut negative for the Denison Trough dawsonites (Fig. 5Band C). Lowest 87Sr/86Sr and d18O values in the GunnedahBasin carbonates are either due to the contribution frommagmatic hydrothermal fluids, or interaction of meteoricwaters with mantle-derived alkaline intrusive rocks. Car-bonates with higher 87Sr/86Sr and d18O values, on the otherhand, reflect interaction of meteoric waters with Permianrocks of the Gunnedah Basin consisting of volcanogenic
clastic rocks with some marine limestone layers (Tadros,1993). Oxygen and strontium isotope compositions ofdawsonites from the Denison Trough represent two differ-ent lithological end members. Fluids that have interactedlargely with Permian clastic rocks containing Rb-rich meta-morphic components at higher temperatures will precipitatedawsonite with higher 87Sr/86Sr but lower d18O values. Bycontrast, relatively lower 87Sr/86Sr and higher d18O valuesof the dawsonites indicate precipitation from fluids thathad interacted dominantly with marine carbonate compo-nents in the Permian succession of the Denison Trough (Ba-ker et al., 1995).
3.3.3. Sm/Nd isotope systematics for carbonates
Three of the six carbonates display eNd(0) values in anarrow range from +3.9 to +5.1 units, indicating thatthe fluids had Nd isotopic compositions slightly moreradiogenic than the chondritic reference CHUR. Carbon-ate sample BS718.10 displays an eNd(0) value of �1.9,being slightly less radiogenic than CHUR but still consid-erably more radiogenic than continental crust at �9 units(here approximated by the “Young Upper Crust” of Kra-mers and Tolstikhin, 1997). Carbonate samplesBOH931.60 and BOH935.20 have highly unusual Nd iso-topic compositions [eNd(0)�45–47], requiring derivationfrom a source that had evolved to extremely radiogeniccompositions at the time of mineral precipitation. Theirisotopic compositions do not correlate with Nd concentra-tions, and plot well outside the broad array displayed bymantle-derived rocks in the 143Nd/144Nd–87Sr/86Sr space(Fig. 12B). With the exception of BOH931.60 andBS718.10, the Sm–Nd isotopic compositions for theremaining carbonates are similar to CHUR.
It is worth noting that two samples BOH931.60 andBOH935.20 have extremely high eNd(0) values of +46–47or eNd(85 Ma) values of +41–46, never reported in any lit-erature for terrestrial rocks. Such values are considerably
Table 6Rb–Sr data for the untreated, leachates, and acid-leached (residues) clay fractions from the Gunnedah Basin.
Sample Host rock Rb (ppm) Sr (ppm) 87Rb/87Sr 87Sr/86Sr ± 2r
5460 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
higher than that of the Depleted MORB Mantle (DMM,�+10), which is considered as the most radiogenic end-member in the mantle source of oceanic basalts (Zindlerand Hart, 1986). Hence, variable contributions from
mantle-derived and crustal sources as the sole sources forNd can be ruled out. This is in contrast to other sampleswith eNd(0) values between +4 and �2, which are consistentwith derivation from alkaline igneous and crustal-like mate-rial as discussed above. Replicate analyses of sampleBOH935.20 give similar 143Nd/144Nd values(0.514988 ± 0.00006 vs. 0.514958 ± 0.00006), excludingthe possibility of contamination as a potential cause. Infact, it is impossible to have a contaminant with such an ex-tremely unusual 143Nd/144Nd value, as typical contami-nants in the laboratory would have 143Nd/144Nd ratiosapproximating the average upper crustal value of 0.5120.Except for these two samples, the most radiogenic samplesever processed in the University of Queensland laboratoryin the past have 143Nd/144Nd ratios <0.5132. In order togenerate eNd(0) values of +46–47 or eNd (85 Ma) valuesof +41–46, the source material must have had an extremelyfractionated REE pattern with a very high Sm/Nd ratio,and this REE-fractionated source must have had a consid-erable crustal residence time to allow radiogenic 143Nd toaccumulate. Sample BOH931.60 displays a strongly frac-tionated REE pattern with a steep kink between Sm andNd (Fig. 7A). The corresponding 147Sm/144Nd = 0.7966 isfour times higher than the chondritic value of 0.1967. If amantle- or crust-derived source material with such a highSm/Nd had a crustal residence time >500 Ma, then it is pos-sible for this material to evolve to the observed eNd(0) val-ues in the two samples. However, sample BOH935.2 onlyhas sub-chondritic 147Nd/144Nd values of 0.1922 to0.2093. Thus, this sample must have experienced a sec-ond-stage REE fractionation (with LREE enrichment) atthe time of carbonate formation, possible around 85 Maago. Hence, purely based on the isotopic signatures of thetwo samples, we propose a two-stage evolution model fortheir origin, i.e., a LREE depletion event resulting in asource material with very high 147Sm/144Nd (e.g., 0.8) oc-curred more than 500 Ma ago, followed by carbonate for-mation through fluid activity around 85 Ma, resulting incarbonates with a large range of 147Sm/144Nd.
Strongly fractionated REE patterns have been observedpreviously in fluorite-rich mineralisation and carbonatecleat fill in coals of the Bowen Basin (Chesley et al., 1991;Uysal et al., 2007b), which were successfully dated by theSm–Nd isochron method owing to the large variations inSm/Nd ratios. Sm/Nd can also be fractionated by co-pre-cipitation of a LREE-rich mineral, such as monazite.
Strongly radiogenic Nd isotopic compositions for sam-ples BOH931.60 and BOH935.20 are not matched bystrongly unradiogenic 87Sr/86Sr ratios (Table 2), indicatingthat both isotope systems were significantly decoupled dur-ing the process of fluid-rock interaction or fluid-migrationand equilibration within the aquifer. Decoupling of Nd,Sr and Pb has been observed previously for hydrothermalmineral deposits, where large variations in Sr and Pb iso-tope ratios were found to be inconsistent with derivationfrom predominant and likely source rocks, such as lime-stones (Bau et al., 2003). Plausible mechanisms for decou-pling of Nd and Sr in the present study includeincongruent trace element leaching from source rocks asproposed by Bau et al. (2003) and different mobility
Fig. 11. Rb–Sr isochron diagrams for leachate-residue-untreatedfractions from the <0.2 lm and 0.5–0.2 size fractions for samplesBOH994.5 and BWN831 (A) and (B), leachate-residue-untreatedfractions from 0.2 lm and 2–0.2 lm size fractions for the sampleBOH939.3 (C), and for leachate-untreated fractions fromthe < 0.2 lm size fraction for sample BS632 (D).
CO2 degassing and trapping in eastern Australia 5461
Author's personal copy
behaviour of Sr and Nd under the prevalent hydrothermalconditions. Hence, isotopic criteria cannot be strictly ap-plied for fingerprinting contributions from bulk sourcerocks, in particular in view of anomalous isotope signaturesas demonstrated for calcite from the Gunnedah Basin.
3.3.4. Illitic clays
Correlations in Fig. 10 indicate that the illite precipitatingfluids had variable origins. Based on the REE content and Euanomaly of illitic clay minerals, two different fluid sourcescan be identified: (1) Meteoric waters that have mixed withmagmatic hydrothermal fluids related to alkali magmatismand/or interacted with alkalic rocks, which are characterisedby high total REE and LREE concentration and no Euanomaly, and (2) meteoric fluids that interacted with rocksof average upper continental crust composition. Moreover,there is an inverse correlation between the d18O values ofthe clays and (Eu/Eu*)PAAS, total REE abundance, andLREE and MREE enrichment in comparison to HREE. Thismay suggest that the less the contribution from fluids thatinteracted mainly with crustal rocks, the more 18O depletedthe illitic clays are that precipitated mainly from the meteoricand magmatic fluid mixture.
Rb–Sr isochron diagrams in Fig. 11 can also provide in-sight into the Sr isotopic compositions of fluids from whichillitic clays have precipitated. Acid leachable components ofthe clay fractions are derived mainly from coexisting fine-grained carbonate minerals. The well-correlated isochrondiagrams in Fig. 11 (data points for leachates, residue anduntreated clays) indicate that Rb–Sr isotopic equilibriumwas achieved between authigenic illite–smectite and carbon-ate mineral phases during hydrothermal events. The lessradiogenic initial 87Sr/86Sr ratio of 0.7056 for BOH 994.5and BWN 831.2 samples is consistent with the less radio-genic 87Sr/86Sr ratios of carbonate vein samples taken fromadjacent locations (Table 2). The initial 87Sr/86Sr value ofclays and the associated carbonates confirms the significantcontribution of mantle-sourced fluids to clay and carbonatemineral precipitation at �94 Ma. Fluids involved in theclay precipitation at �85 Ma had a more radiogenic signa-ture indicating significant interaction of fluids with moreevolved crustal rocks. Particularly, the initial 87Sr/86Sr va-lue of 0.7276 for sample BOH 939.3 requires significant
Sr input from highly evolved Rb-rich crustal rocks. Themost likely candidate for such source rocks is the mica-richmetasedimentary basement of the Gunnedah Basin(Lachlan Fold Belt; Leith, 1993). The initial 87Sr/86Sr value
Fig. 12. (A): eNd vs time (Myr) illustrating the Nd isotopicevolution of Gunnedah Basin carbonates with reference toDepleted Mantle, Chondritic Uniform Reservoir (CHUR) andContinental Crust (Kramers and Tolstikhin, 1997). Note the highlyunusual Nd isotopic composition of two carbonates samples witheNd > 40 units, requiring multiple-stage fractionation events. (B)143Nd/144Nd vs 87Sr/86Sr isotope ratio diagram showing Gunnedahcarbonates in comparison to the globally observed mantle–crustarray (schematic). Note that the Gunnedah carbonates have Ndisotopic compositions comparable to the Depleted Mantle but atconsiderably more radiogenic 87Sr/86Sr, indicative of decoupling ofboth isotope systems in the past.
Table 7Sm–Nd isotope data for carbonates from the Gunnedah Basin.
5462 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
of 0.7091 for sample BS2C 632 can be explained by signif-icant Sr contribution either from marine limestone layers orvolcanigenic Permian clastic rocks of the Gunnedah Basin.All three different Sr initial values for the clays are similarto 87Sr/86Sr values of the carbonate vein samples takenfrom close proximity in the same holes. Different depthsof borehole BOH (e.g., 931–939.3 m and 994.5–998.5 m)represent different precipitation ages with different Sr iso-tope values of clays and carbonates. This implies that clayand carbonate generations developed as a result of chan-nelized hydrothermal fluid circulation events (rather thanpervasive and uniform fluid flow) in response to Gondwanarifting and the associated magmatism.
3.4. CO2 sources
The range of carbon isotope compositions of calcites andankerites from the Gunnedah and Bowen Basins is highlyvariable. We previously interpreted the negative d13C valuesfor the Bowen Basin samples to reflect CO2 derived fromthermal degradation of organic matter and attributed themore positive d13C values to exchange reactions in a mixedCO2–CH4 bearing fluid (Carothers and Kharaka, 1980; Uy-sal et al., 2000a). Based on petrographic observations, sider-ite from the Gunnedah Basin, similarly to the siderite fromthe Bowen Basin, seems to have formed as an early diageneticphase under reducing conditions at lower temperatures(<80 �C). Carbon and oxygen isotope compositions of theGunnedah Basin siderites exhibit a similar trend to the Bo-wen Basin siderites, with both positive and negative d13C val-ues indicative of CO2 derived from organic sources andmethanogenesis, respectively (cf., Uysal et al., 2000a).
We previously interpreted the overlapping d13C valuesof Gunnedah and Bowen Basin carbonates to representsimilar CO2 sources (Fig. 5A). However, despite the similard13C ranges, there are marine components in the stratigra-phy of the Gunnedah Basin, which may provide additionalpotential CO2 sources in this basin. Indeed, uniform87Sr/86Sr ratios for the dawsonite are consistent with mar-ine Sr isotope compositions. Although Sr and other cationsmay be of marine origin, CO2 for the dawsonite formationmay have been derived largely from different sources. It islikely that magmatic CO2 has largely contributed to theprecipitation of the dawsonite in view of the narrow rangeof d13C values across the BGS basin system (cf., Bakeret al., 1995). In summary, possible CO2 sources for theinvestigated carbonates include (1) thermal degradation oforganic matter (calcite and ankerite), (2) bacterial methano-genesis (siderite), and (3) magmatic gas and/or marine car-bonate (dawsonite and some calcite–ankerite). Accordingto comprehensive carbon isotope data from widespreadnatural gas fields, gases with highest CO2 contents are de-rived from inorganic sources such as magmatism or marinecarbonates (Boreham et al., 2001).
3.5. Implications for CO2 mineral trapping and adsorption in
coal systems
There seems to be a correlation among the intensity ofcalcite–ankerite vein/cement formation, dawsonite veining
and rock lithology in the Gunnedah Basin coal measures.Early diagenetic siderite and kaolinite in the claystonesare abundant between the Hoskisson coal seams and intheir immediate roof and floor sections. Remarkably, daw-sonite veins occur usually within such impermeable litho-logic units, whereas abundant calcite–ankerite veins andcements are observed mostly in more permeable units asso-ciated with sandstone-conglomerate and volcaniclasticrocks. In some cases the precipitation of early diageneticsiderite and kaolinite in finer grained units may have re-duced the permeability sufficiently to prevent the circula-tion of hydrothermal fluids from which the illitic claysand calcite–ankerite precipitated in more permeable zones.CO2 of magmatic/mixed origin has probably reached everystratigraphic level in the Gunnedah Basin, but it seems tohave dissipated as carbonate precipitation in the more per-meable and cation-rich rock units. By contrast, as CO2 en-tered the Hoskisson seams, mineral reaction was inhibited,most likely due to reduced fluid flow and limited cationavailability. Increasing accumulation and oversaturationwith CO2 would have resulted in high CO2 gas pressure thatpropagated along cleats and fractures. Under such condi-tions, because of the low permeability of the system andhigh sorption capacity of CO2 on coal (White et al.,2005a,b), much of the CO2 has remained adsorbed in thecoal seams since the Cretaceous.
Several authors have argued from studies of CO2-richgas accumulations that the role of mineral reactions in trap-ping of CO2 is limited, and that dissolution in formationwaters is the major sink for CO2 (Gilfillan et al., 2009; Wil-kinson et al., 2009). The relative importance of dawsoniteformation in mineral trapping at engineered CO2 storagesites is also controversial because dawsonite is a relativelyrare authigenic mineral in sedimentary basins despite mod-elling studies that predict its formation at the expense offeldspar in high CO2 environments (e.g., White et al.,2005a,b; Xu et al., 2007). This could be because many res-ervoir sandstones lack reactive phases such as sodic plagio-clase and volcanic rock fragments, which is not the case inthe BGS basin system where dawsonite occurs in the coalmeasures and also in sandstones and siltstones in both mar-ine and nonmarine facies (Baker et al., 1995). The similarcarbon isotope composition of dawsonites across the BGSbasin system and common association with CO2-rich natu-ral and coal seam gas accumulations suggest that dawsoniteformed as a result of influx of magmatic CO2 most likelyunder open system conditions (cf., Baker et al., 1995). Thisis supported by carbon isotope studies that confirm themagmatic origin of the CO2 in CO2-rich natural and coalseam gases in the BGS basin system (Draper and Boreham,2006; Faiz et al., 2007). Dawsonite precipitation occurred inresponse to elevated CO2 fugacity and alkalinity in the geo-chemically more evolved environment at the latest stage ofthe fluid flow process, with cations derived from the disso-lution of feldspars and clay minerals.
Our findings indicate that dawsonites in reservoir andcaprock formations of the BGS basin system formed as aresult of magmatic/mantle CO2 influx and have remainedstable over many 10s of millions of years. The CO2 mineraltrapping capacity depends on primary mineral composition
CO2 degassing and trapping in eastern Australia 5463
Author's personal copy
and permeability. In this context dawsonite in the coalseams is essentially a low volume fracture fill, whereas daw-sonite in the sandstones and siltstones fills pores and frac-tures and also replaces framework grains reaching localabundances exceeding 10%. More broadly, dawsonite min-eralization in basinal settings worldwide such as the Perm-ian Groeden Sandstone (Italy) and Triassic Lam Formation(Yemen) is thought to be related to deeply sourced mag-matic/mantle CO2 that migrated upwards along fracturesand more permeable zones (Wopfner and Hocker, 1987;Worden, 2006). By contrast, dawsonite is absent in lithicsandstones of the Pretty Hill Formation (South Australia)where ferroan carbonates formed at the expense of lithicfragments in response to high partial pressures of magmaticCO2 in the reservoir (Watson et al., 2004). This likely re-flects variations in the carbonate buffering capacity of thelatter formation that may have been insufficient to producedawsonite saturated conditions at the P–T conditions in thereservoir.
4. SUMMARY AND CONCLUSIONS
We conducted clay mineralogical, stable and radiogenicisotope, and trace element analyses of the carbonate andclay minerals from clastic sedimentary rocks of the Gunne-dah Basin and the Denison Trough in the Bowen Basin.Our work shows that carbonate (calcite, ankerite, sideriteand dawsonite) and clay (kaolinite and illite–smectite) min-eral precipitation occurred intensively in the Permo-TriassicGunnedah Basin as a result of Mesozoic regional fluid flowevents in a CO2-rich environment.
The distribution of illitic clay types in the Gunnedah Ba-sin holes is not depth-related and changes irregularly withinrelatively thin stratigraphic intervals. This is attributed tochanges in fluid composition, temperature, permeabilityand fluid/rock ratio in a hydrothermal system. As inferredfrom Rb–Sr isochron age data of the illitic clays, the hydro-thermal process in the Gunnedah Basin occurred as epi-sodic cycles at �85 Ma and �95 Ma in association withGondwana rifting accompanied by alkaline magmatism ineastern Australia. Stable isotope data of carbonate and clayminerals are indicative of meteoric waters from a high-lati-tude environment as the main fluid source, whereas traceelement, Sr and Nd isotope data highlight that the fluidsalso had magmatic and/or crustal components, with a pos-sible input from marine carbonates for some samples. TheGunnedah Basin carbonates have superchondritic Zr/Hfratios that correlate positively with Y/Ho ratios, with thelatter ratios increasing with decreasing 87Sr/86Sr values.These correlations and the positive correlation between87Sr/86Sr and d18O values are interpreted as fluid contribu-tions from both mantle and crustal end members, such asmagmatic hydrothermal fluids and meteoric waters thathad interacted with mantle-derived alkaline intrusive rocksor clastic and marine carbonate rocks of the GunnedahBasin.
Oxygen and strontium isotope ratios of dawsonite ce-ments from the Denison Trough sandstones represent twodifferent lithological end members. Dawsonites with higher87Sr/86Sr but lower d18O values reflect fluid interaction with
Permian clastic rocks, whereas lower 87Sr/86Sr and higherd18O values of the dawsonites are interpreted to reflect in-creased contribution from fluids that had interacted domi-nantly with marine carbonates in the Permian successionof the Denison Trough. Based on carbon isotope data,CO2 used for calcite and ankerite precipitation originatedmainly from thermal degradation of organic matter andmagmatic gas, whereas the CO2 consumed for the dawson-ite formation is inferred to have derived from magmaticand marine sources. In the low permeability environments(particularly in coal seams), the increasing accumulationand oversaturation of CO2 promote the precipitation ofdawsonite.
ACKNOWLEDGMENTS
The authors thank the CO2CRC for sponsoring this researchand acknowledge the funding provided by the Commonwealth ofAustralia through the CRC Program. We greatly acknowledgeJoan Esterle for her encouragement to conduct this research andher advice and support for the sampling. We also thank our indus-try collaborators for permission to sample core and the QueenslandMuseum for access to catalogued material from the DenisonTrough. The paper was significantly improved by the constructiveand detailed reviews of Fred Longstaffe and an anonymous re-viewer, who are gratefully acknowledged. We thank Associate Edi-tor, Edward Ripley for his helpful comments and editorialhandling.
REFERENCES
Bachu S. (2000) Sequestration of CO2 in geological media: criteriaand approach for site selection in response to climate change.Energy Convers. Manage. 41, 953–970.
Baker J. C., Bai G. P., Hamilton P. J., Golding S. D. and Keene J.B. (1995) Continental-scale magmatic carbon-dioxide seepagerecorded by dawsonite in the Bowen–Gunnedah–Sydney BasinSystem, Eastern Australia. J. Sediment. Res. A 65, 522–530.
Barclay S. A. and Worden R. H. (2000) Geochemical modelling ofdiagenetic reactions in a sub-arkosic sandstone. Clay Miner. 35,
57–67.
Battaglia S., Gherardi F., Gianelli G., Leoni L. and Origlia F.(2007) Clay mineral reactions in an active geothermal area (Mt.Amiata, southern Tuscany, Italy). Clay Miner. 42, 353–372.
Bau M., Romer R. L., Luders V. and Dulski P. (2003) Tracingelement sources of hydrothermal mineral deposits: REE and Ydistribution and Sr–Nd–Pb isotopes in fluorite from MVTdeposits in the Pennine Orefield, England. Miner. Deposita 38,
992–1008.
Bau M. (1996) Controls on the fractionation of isovalent traceelements in magmatic and aqueous systems: Evidence from Y/Ho, Zr/Hf, and lanthanide tetrad effect. Contrib. Mineral.
Petrol. 123, 323–333.
Bau M. and Dulski P. (1995) Comparative-study of yttrium andrare-earth element behaviors in fluorine-rich hydrothermalfluids. Contrib. Mineral. Petrol. 119, 213–223.
Boles J. R. (2008) Authigenic minerals, in Access Science,�McGraw-Hill Companies, 2008. <http://www.accessscience.com>.
Boreham C. J., Hope J. M. and Hartung-Kagi B. (2001)Understanding source, distribution and preservation of Aus-tralian natural gas: a geochemical perspective. APPEA J. 41,
523–547.
5464 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
Author's personal copy
Carothers W. W. and Kharaka Y. K. (1980) Stable carbon isotopesof HCO-3 in oil-field waters-implications for the origin of CO2.Geochim. Cosmochim. Acta 44, 323–332.
Carothers W. W., Lanford H. A. and Rosenbauer R. J. (1988)Experimental oxygen isotope fractionation between siderite–water and phosphoric acid liberated CO2-siderite. Geochim.
Cosmochim. Acta 52, 2445–2450.
Chesley J. T., Halliday A. N. and Scrivener R. C. (1991)Samarium–neodymium direct dating of fluorite mineralization.Science 252, 949–951.
Chiodini G., Baldini A., Barberi F., Carapezza M. L., CardelliniC., Frondini F., Granieri D. and Ranaldi M. (2007) Carbondioxide degassing at Latera caldera (Italy): evidence ofgeothermal reservoir and evaluation of its potential energy. J.
Geophys. Res. Solid Earth 112(B12), 17.
Chiodini G., Cardellini C., Amato A., Boschi E., Caliro S.,Frondini F. and Ventura G. (2004) Carbon dioxide Earthdegassing and seismogenesis in central and southern Italy.Geophys. Res. Lett. 31, L07615.
Clauer N. and Chaudhuri S. (1995). , p. 359.Clauer N., Chaudhuri S., Kralik M. and Bonnotcourtois C. (1993)
Effects of experimental leaching on Rb–Sr and K–Ar isotopicsystems and REE contents of diagenetic illite. Chem. Geol. 103,
1–16.
Clayton R. N., Friedman I., Graf D. L., Mayeda T. K., Meents W.F. and Shimp N. F. (1966) The origin of saline formationwaters. J. Geophys. Res. Solid Earth 71, 3869–3882.
Draper J. J. and Boreham C. J. (2006) Geological controls onexploitable coal seam gas distribution in Queensland. APPEA
J. 46, 343–366.
Eggins S. M., Woodhead J. D., Kinsley L. P. J., Mortimer G. E.,Sylvester P., McCulloch M. T., Hergt J. M. and Handler M. R.(1997) A simple method for the precise determination of >=40trace elements in geological samples by ICPMS using enrichedisotope internal standardisation. Chem. Geol. 134, 311–326.
Elliott L. (1993) Post-carboniferous tectonic evolution of easternAustralia. APPEA J. 33, 215–236.
Faiz M. M., Saghafi A., Barclay S. A., Stalker L., Sherwood N. R.and Whitford D. J. (2007) Evaluating geological sequestrationof CO2 in bituminous coals: The southern Sydney Basin,Australia as a natural analogue. Int. J. Greenhouse Gas Control
1, 223–235.
Gilfillan S. M. V., Lollar B. S., Holland G., Blagburn D., StevensS., Schoell M., Cassidy M., Ding Z., Zhou Z., Lacrampe-Couloume G. and Ballentine C. J. (2009) Solubility trapping information water as dominant CO2 sink in natural gas fields.Nature 458, 614–618.
Golab A. N., Carr P. F. and Palamara D. R. (2006) Influence oflocalised igneous activity on cleat dawsonite formation in LatePermian coal measures, Upper Hunter Valley, Australia. Int. J.
Coal Geol. 66, 296–304.
Goldstein S. L., Onions R. K. and Hamilton P. J. (1984) A Sm–Ndisotopic study of atmospheric dusts and particulates from majorriver systems. Earth Planet. Sci. Lett. 70, 221–236.
Hower J., Eslinger E. V., Hower M. E. and Perry E. A. (1976)Mechanism of burial metamorphism of argillaceous sediment.1. Mineralogical and chemical evidence. Geol. Soc. Am. Bull. 87,
725–737.
Huang W. L. and Longo J. M. (1994) Experimental studies ofsilicate-carbonate reactions. 2. Applications to steam floodingof oil sands. Appl. Geochem. 9, 523–532.
Jacobsen S. B. and Wasserburg G. J. (1980) Sm–Nd isotopicevolution of chondrites. Earth Planet. Sci. Lett. 50, 139–155.
Jennings S. and Thompson G. R. (1986) Diagenesis of Pleistocenesediments of the Colorado River-Delta, Southern-California. J.
Sediment. Petrol. 56, 89–98.
Kamber B. S., Greig A. and Collerson K. D. (2005) A new estimatefor the composition of weathered young upper continental crustfrom alluvial sediments, Queensland, Australia. Geochim.
Cosmochim. Acta 69, 1041–1058.
Kaszuba J. P., Williams L. L., Janecky D. R., Hollis W. K. andTsimpanogiannis I. N. (2006) Immiscible CO2–H2O fluids in theshallow crust. Geochem. Geophys. Geosyst. 7. doi:10.1029/
2005gc001107.
Kramers J. D. and Tolstikhin I. N. (1997) Two terrestrial leadisotope paradoxes, forward transport modelling, core forma-tion and the history of the continental crust. Chem. Geol. 139,
75–110.
Leith E. C. (1993) The floor of the Gunnedah Basin north of theLiverpool Range. In The Gunnedah Basin, New South Wales
(ed. N.Z. Tadros). Geological Survey of New South Wales,Memor 12, 335–348.
Longstaffe F. J. and Ayalon A. (1990) Hydrogen-isotope geochem-istry of diagenetic clay-minerals from Cretaceous sandstones,Alberta, Canada – Evidence for exchange. Appl. Geochem. 5,
657–668.
Ludwig K. R. (2003) User’s Manual for Isoplot/Ex Version2.02, A Geochronological Toolkit for Microsoft ExcelBerkeley Geochronology Center Spec. Pub. 1a, Berkeley, CA,USA.
McCrea J. M. (1950) On the isotopic chemistry of carbonates and apaleotemperature scale. J. Chem. Phys. 18, 49–857.
McCulloch M. T. and Wasserburg G. J. (1978) Sm–Nd and Rb–Srchronology of continental crust formation. Science 200, 1003–
1011.
Moore D. M. and Reynolds R. C. (1989). , p. 322.Moore J., Adams M., Allis R., Lutz S. and Rauzi S. (2005)
Mineralogical and geochemical consequences of the long-termpresence of CO2 in natural reservoirs: an example from theSpringerville–St. Johns Field, Arizona, and New Mexico, USA.Chem. Geol. 217, 365–385.
Mutlu H., Gulec� N. and Hilton D. R. (2008) Helium–carbonrelationships in geothermal fluids of western Anatolia, Turkey.Chem. Geol. 247, 305–321.
Nesbitt B. E., Muehlenbachs K. and Murowchick J. B. (1989)Genetic implications of stable isotope characteristics of meso-thermal Au deposits and related Sb and Hg deposits in theCanadian Cordillera. Econ. Geol. 84, 1489–1506.
O’Neil J. R., Clayton R. N. and Mayeda T. K. (1969) Oxygenisotope fractionation in divalent metal carbonates. J. Chem.
Phys. 51, 5547–5548.
O’Sullivan P. B., Kohn B. P., Foster D. A. and Gleadow A. J. W.(1995) Fission track data from the Bathurst Batholith: evidencefor rapid mid-Cretaceous uplift and erosion within the easternhighlands of Australia. Aust. J. Earth Sci. 42, 597–607.
Pollastro R. M. (1993) Considerations and applications of theillite–smectite geothermometer in hydrocarbon-bearing rocks ofMiocene to Mississippian age. Clay Clay Miner. 41, 119–133.
Rosenbaum J. and Sheppard S. M. F. (1986) An isotopic study ofsiderites, dolomites and ankerites at high temperature. Geochim.
Cosmochim. Acta 50, 1147–1150.
Salvi S. and Williamsjones A. E. (1990) The role of hydrothermalprocesses in the granite-hosted Zr, Y, REE deposit at StrangeLake, Quebec Labrador – Evidence from fluid inclusions.Geochim. Cosmochim. Acta 54, 2403–2418.
Salvi S. and Williamsjones A. E. (1996) The role of hydrothermalprocesses in concentrating high-field strength elements in theStrange Lake peralkaline complex, northeastern Canada. Geo-
chim. Cosmochim. Acta 60, 1917–1932.
Sharma T. and Clayton R. N. (1965) Measurement of 18O/16Oratios of total oxygen in carbonates. Geochim. Cosmochim. Acta
29, 1347–1353.
CO2 degassing and trapping in eastern Australia 5465
Author's personal copy
Sharp Z. D. (1990) Laser-based microanalytical method for thein situ determination of oxygen isotope ratios of silicates andoxides. Geochim. Cosmochim. Acta 54, 1353–1357.
Sheppard S. M. F. and Schwarcz H. P. (1970) Fractionation ofcarbon and oxygen isotopes and magnesium between coexitingmetamorphic calcite and dolomite. Contrib. Mineral. Petrol. 26,
161–198.
Sheppard S. M. F. (1986) Characterization and isotopic variationsin natural waters. In Stable Isotopes in High Temperature
Geological Processes: Reviews in Mineralogy (eds. J. W. Valley,H. P. Taylor, Jr., J. R. O’Neil) 16, pp. 165–183.
Sheppard S. M. F. and Gilg H. A. (1996) Stable isotopegeochemistry of clay minerals. Clay Mineral. 31, 1–24.
Tadros N. Z. (1993) Lithostratigraphy. In The Gunnedah Basin,
New South Wales (ed. N. Z. Tadros). Geological Survey of NewSouth Wales, Memor 12, 95–134.
Taylor H. P. (1990) Oxygen and hydrogen isotope constrains onthe deep circulation of surface waters into zones of hydrother-mal metamorphism and melting. In: The Role of Fluids in
Crustal Processes; Studies in Geophysics, vol. NationalResearch Council, National Academy Press, Washington,DC, pp. 72–95.
Taylor S. R. and McLennan S. M. (1985) The Continental Crust: its
Composition and Evolution. Blackwell, Oxford, 312 pp.Truesdell A. H. and Hulston J. R. (1980) Isotopic evidence on
environments of geothermal systems. In Handbook of Environ-
mental Isotope Geochemistry 1 (eds. P. Fritz and J. C. Fontes).Elsevier, Amsterdam–Oxford–New York, pp. 179–226.
Uysal I. T., Feng Y., Zhao J. X., Altunel E., Weatherley D.,Karabacak V., Cengiz O., Golding S. D., Lawrence M. G. andCollerson K. D. (2007a) U-series dating and geochemicaltracing of late Quaternary travertine in co-seismic fissures.Earth Planet. Sci. Lett. 257, 450–462.
Uysal I. T., Zhao J. X., Golding S. D., Lawrence M. G., GliksonM. and Collerson K. D. (2007b) Sm–Nd dating and rare-earthelement tracing of calcite: Implications for fluid-flow events inthe Bowen Basin, Australia. Chem. Geol. 238, 63–71.
Uysal I. T., Golding S. D. and Glikson M. (2000a) Petrographicand isotope constraints on the origin of authigenic carbonateminerals and the associated fluid evolution in Late Permian coalmeasures, Bowen Basin (Queensland), Australia. Sediment.
Geol. 136, 189–206.
Uysal I. T., Golding S. D. and Baublys K. (2000b) Stable isotopegeochemistry of authigenic clay minerals from Late Permiancoal measures, Queensland, Australia: implications for theevolution of the Bowen Basin. Earth Planet. Sci. Lett. 180, 149–
162.
Uysal I. T., Golding S. D. and Audsley F. (2000c) Clay-mineralauthigenesis in the Late Permian coal measures, Bowen Basin,Queensland, Australia. Clay Clay Miner. 48, 351–365.
Uysal I. T. and Golding S. D. (2003) Rare earth elementfractionation in authigenic illite–smectite from Late Permianelastic rocks, Bowen Basin, Australia: implications for physico-
chemical environments of fluids during illitization. Chem. Geol.
193, 167–179.
Uysal I. T., Feng Y. X., Zhao J. X., Isik V., Nuriel P. and GoldingS. D. (2009) Hydrothermal CO2 degassing in seismically activezones during the late Quaternary. Chem. Geol. 265, 442–454.
Veevers J. J. (1984). , p 418.Veevers J. J., Powell C. M. and Roots S. R. (1991) Review of sea-
floor spreading around Australia. 1. Synthesis of the patterns ofspreading. Aust. J. Earth Sci. 38, 373–389.
Waight T. E., Weaver S. D., Maas R. and Eby G. N. (1998) FrenchCreek Granite and Hohonu Dyke Swarm, South Island, NewZealand: Late Cretaceous alkaline magmatism and the openingof the Tasman Sea. Aust. J. Earth Sci. 45, 823–835.
Watson M. N., Zwingmann N. and Lemon N. M. (2004) TheLadbroke Grove–Katnook carbon dioxide natural laboratory:a recent CO2 accumulation in a lithic sandstone reservoir.Energy 29, 1457–1466.
White C. M., Smith D. H., Jones K. L., Goodman A. L., Jikich S.A., LaCount R. B., DuBose S. B., Ozdemir E., Morsi B. I. andSchroeder K. T. (2005a) Sequestration of carbon dioxide in coalwith enhanced coalbed methane recovery – a review. Energy
Fuels 19, 659–724.
White S. P., Allis R. G., Moore J., Chidsey T., Morgan C., GwynnW. and Adams M. (2005b) Simulation of reactive transport ofinjected CO2 on the Colorado platea, Utah, USA. Chem. Geol.
217, 387–405.
Wilkinson M., Haszeldine R. S., Fallick A. E., Odling N., Stoker S.J. and Gatliff R. W. (2009) CO2-mineral reaction in a naturalanalogue for CO2 storage – implications for modelling. J.
Sediment. Res. 79, 486–494.
Wood S. A. (1990) The aqueous geochemistry of the rare-earthelements and yttrium. 2. Theoretical predictions of speciation inhydrothermal solutions to 350 �C at saturation water–vaporpressure. Chem. Geol. 88, 99–125.
Wopfner H. and Hocker C. F. W. (1987) The Permian GroedenSandstone between Bozen and meran (northern Italy), a habitatof dawsonite and nordstrandite. Neues Jahrbuch fur Geologie
and Palaontologie, Monatshefte, 161–176.Worden R. H. (2006) Dawsonite cement in the Triassic Lam
Formation, Shabwa Basin, Yemen: a natural analogue for apotential mineral product of subsurface CO2 storage forgreenhouse gas reduction. Mar. Petrol. Geol. 23, 61–77.
Yeh H. W. (1980) D/H ratios and late-stage dehydration of shalesduring burial. Geochim. Cosmochim. Acta 44, 341–352.
Xu T., Apps J. A., Pruess K. and Yamamoto H. (2007) Numericalmodelling of injection and mineral trapping of CO2 with H2Sand SO2 in a sandstone formation. Chem. Geol. 242, 319–346.
Zindler A. and Hart S. (1986) Chemical geodynamics. Annu. Rev.
Earth Planet. Sci. 14, 493–571.
Associate editor: Edward M. Ripley
5466 I.T. Uysal et al. / Geochimica et Cosmochimica Acta 75 (2011) 5444–5466
