Permo-Pennsylvanian Palaeotemperatures from Fe-Oxide and Phyllosilicate Δ18o Values ⁎ Neil J

Earth and Planetary Science Letters 253 (2007) 159–171 www.elsevier.com/locate/epsl

Permo-Pennsylvanian palaeotemperatures from Fe-Oxide and phyllosilicate δ18O values ⁎ Neil J. Tabor

Department of Geological Sciences, Southern Methodist University, Dallas, TX, 75275-0395, United States Received 13 February 2006; received in revised form 9 October 2006; accepted 11 October 2006 Available online 21 November 2006 Editor: H. Elderfield

Abstract

The oxygen isotope composition of fossil roots that have been permineralized by hematite are presented from eight different stratigraphic levels spanning the Upper Pennsylvanian and Lower Permian strata of north-central Texas. Hematite δ18O values range from −0.4% to 3.7%. The most negative δ18O values occur in the upper Pennsylvanian strata, and there is a progressive trend toward more positive δ18O values upward through the lower Permian strata. This stratigraphic pattern is similar in magnitude and style to δ18O values reported for penecontemporaneous authigenic palaeosol phyllosilicates and calcites, suggesting that all three minerals record similar paragenetic histories that are probably attributed to temporal palaeoenvironmental changes across the Late Pennsylvanian and Early Permian landscapes. Palaeotemperature estimates based on paired δ18O values between penecontemporaneous hematite and phyllosilicate samples suggest these minerals co-precipitated at relatively low temperatures that are consistent with a supergene origin in a low-latitude soil-forming environment. Hematite–phyllosilicate δ18O pairs indicate (1) relatively low soil temperatures (∼24±3 °C) during deposition of the upper Pennsylvanian strata followed by (2) a considerable rise in soil temperatures (∼35–37±3 °C) during deposition of the lowermost Permian strata. Significantly, δD and δ18O values of contemporaneous phyllosilicates provide single mineral palaeotemperature estimates that are analytically indistinguishable from temperature estimates based on hematite– phyllosilicate oxygen isotope pairs. The results between the two temperature-proxy methods suggest that the inferred large temperature change across the Upper Pennsylvanian–Lower Permian boundary might be taken seriously. If real, such a significant climate change would have undoubtedly had far-reaching ecological effects within this region of Pangaea. Notably, there are important lithological and palaeobotanical changes, such as disappearance of coal and coal swamp floras, across the Upper Pennsylvanian–Early Permian boundary of north-central Texas that may be consistent with major climatic change toward warmer conditions. © 2006 Elsevier B.V. All rights reserved.

Keywords: Hematite; Phyllosilicate; Oxygen isotope equilibrium; Palaeotemperatures; Permian

1. Introduction mineralogical component of ancient sedimentary strata [1]. Many recent studies have demonstrated that the Fe(III) oxyhydroxides commonly form in near- isotopic composition of Fe(III) oxyhydroxides, as well surface sedimentary environments and are an abundant as other minerals that typically form in low temperature environments such as calcite and phyllosilicate, may ⁎ Tel.: +1 214 768 4175. provide important palaeoenvironmental information E-mail address: [email protected]. related to the conditions of mineral formation in the

Fig. 1. Pangaean Continental Reconstruction and Permo-Pennsylvanian physiography (∼300 Ma BP). The palaeogeographic position of the north- central Texas study area is marked by the square symbol along western Pangaea, south of the equator. Modified from [47]. (B) Stratigraphic and Geological map of the field area in north central Texas. The stratigraphic column (left) shows epoch and stage boundaries according to conventionalisms for West Texas, whereas the thin dashed lines upon the geological map (right) show stage boundaries according to the International Geological Timescale [48] based upon conodont biostratigraphic zonation in north-central Texas strata (B.Wardlaw, pers. comm., 2005). Bold letters A, B, C, D, E, F, G, H mark approximate locations from which Fe-oxides were sampled in this study. See text for discussion. N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 161 early burial environment. In particular, these minerals 2. Geologic background can provide information about the δ18O values of local meteoric water and temperature of crystallization 2.1. Lithostratigraphy [1–20]. In this regard, pedogenic Fe(III) oxides and phyllo- The dominantly terrestrial Upper Pennsylvanian and silicates have the potential to provide isotopic records of Lower Permian succession of north-central Texas was palaeoclimate, and palaeoclimatic change, through a deposited in three broad depositional belts (lower and succession of sedimentary rocks. To date, however, upper coastal plain, and piedmont facies) distributed there has been no systematic stratigraphic analysis of across the low-sloping eastern shelf of the Midland basin. oxygen isotope compositions of co-existing Fe-oxides The succession comprises ∼1100mofupwardfining and phyllosilicates preserved in sedimentary strata fluvio-alluvial cycles [21] and contains abundant, well- [17,20]. developed palaeosols [17,22]. The study area remained in This study presents mineralogical, chemical and the western equatorial region of Pangaea, within 5° of the oxygen isotope compositions of Fe(III) oxides from equator, throughout Permo-Pennsylvanian time [M. Upper Pennsylvanian through Early Permian sedimen- Steiner, unpublished data; 23]. Fe (III) oxides occur as tary strata of the eastern Midland basin, north-central permineralized plant fossils of in-situ root systems within Texas, U.S.A. These new Fe-oxide data are compared mudstones and claystones of overbank-floodplain depos- with previously published oxygen isotope values of its (Fig. 2). Reflected light microscopy of these samples penecontemporaneously formed pedogenic phyllosili- show preservation of original cellular root morphology of cates of the eastern Midland basin to investigate the secondary xylem (Fig. 2). This sort of preservation (1) isotopic models of mineral formation in the early suggests these roots were likely permineralized soon after burial environment, (2) the feasibility of mineral-water death, in the very early burial environment, prior to any oxygen isotope fractionation equations, and (3) Permo- significant humification of the woody root material [24]. Pennsylvanian palaeoclimate over western equatorial Regionally extensive fluvial sandstone sets [21] are Pangaea (Fig. 1A). intercalated with marine limestone marker beds throughout

Fig. 2. (A) Photographic image of permineralized fossil root structures (Sample E). Scale is marked in millimeters. (B) Reflected light image of polished transverse section from sample E showing permineralized celluar structure of the root system. Scale bar is 1 mm. (C) Magnified reflected light image of permineralized cells of the secondary xylem in (B). Dark areas define cell walls, whereas light areas define cellular lumen. Field of view is 1 mm. (D) Close up of C. Light circular area is ∼0.1 mm across. See Text for discussion. 162 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 the succession and define a high-resolution stratigraphic Table 1 framework for correlation throughout the study region. Name, unit, and position of the Fe-oxide samples, and mineralogy as determined by X-ray diffraction Upper Pennsylvanian (and/or Carboniferous) and Lower Permian epoch and stage boundaries have been identified Sample name Stratigraphic Stratigraphic position Mineralogy on the basis of conodont and fusilinid biostratigraphy of unit (m)1 intercalated marine rocks in the north-central Texas strata H Clear Fork Gp. 850 Hematite (B. Wardlaw, pers. comm., 2005; Fig. 1B). The Pennsyl- G Clear Fork Gp. 805 Hematite – F Clear Fork Gp. 760 Hematite vanian Permian boundary [301± 2 Ma; 25] occurs within E Clear Fork Gp. 710 Hematite the upper Markley Formation of the Bowie Group D Nocona Fm. 315 Hematite (Fig. 1B). C Markley Fm. 230 Hematite, Goethite, 2.2. Previous palaeoclimate inferences Kaolinite B Markley Fm. 205 Hematite A Markley Fm. 180 Hematite Palaeosols are an important stratigraphic component of the Upper Pennsylvanian and Lower Permian succession of north-central Texas [22,26]. Tabor and Montañez floras across the Permo-Pennsylvanian landscape [26]. [17,22] defined eight pedotypes [27] in Permo-Pennsyl- Nevertheless, temperature estimates based upon phyllo- vanian strata of north-central Texas based on palaeosol silicate δ18OandδD values are not without uncertainties. macro- and micromorphological characteristics. Those Necessary assumptions that must be made for these studies interpreted the stratigraphic distribution of pedo- palaeotemperature estimates from phyllosilicate include types to record a relatively rapid transition from humid (1) chemical equilibrium with meteoric water during conditions typical of the Late Pennsylvanian to signifi- Permo-Pennsylvanian pedogenesis, (2) knowledge of the cantly drier conditions in the earliest Permian. A climate phyllosilicate-water oxygen isotope fractionation factors trend toward increasingly drier conditions is recorded and (3) little or no post-pedogenic alteration that would throughout the Lower Permian by changes in the strati- have resulted in isotopic change of original δ18OorδD graphic distribution of palaeosol morphologies [22]. values of the phyllosilicates. Without an independent Moreover, the palaeosols proximal to the Pennsylva- method of determining Permo-Pennsylvanian palaeotem- nian–Permian boundary exhibit morphological character- peratures, such as oxygen isotope-pair palaeothermometry istics that likely record rapid onset of seasonality in based on the δ18O values of coexisting palaeopedogenic precipitation coincident with the transition to drier Fe-oxides and phyllosilicate, it is difficult to evaluate the conditions [22,26]. validity of the aforementioned assumptions and the In addition to Fe(III) oxides, certain Permo-Pennsyl- resulting palaeotemperature estimates. vanian-age pedotypes of the Eastern Shelf of the Midland basin also contain pedogenic phyllosilicates and calcites 3. Methods [17–19]. The oxygen and hydrogen isotope compositions of the pedogenic phyllosilicates were used to estimate Eight different Fe-oxide samples were taken from changes in meteoric water δ18Ovaluesfrom−5.5% to the Upper Pennsylvanian Markley Formation and the −3.5% and temperatures from 22±3 °C to 33±3 °C Lower Permian Archer City, Nocona, and Clear Fork through Permo-Pennsylvanian time in this region of Formations (Table 1). Samples were initially collected in western equatorial Pangea [18,19]. Furthermore, the aluminum foil. The chemical pretreatments for the iron- phyllosilicate δDandδ18Odatawereusedtodetermine oxide samples follow the methods of [6]. Samples were that isotope disequilibrium with coexisting calcite that ground in a corundum mortar and pestle under reagent- likely reflects evaporative enrichment of soil water δ18O grade acetone and sized by passage through a 60 μm values prior to calcite precipitation [18]. Collectively, brass sieve. Only powders from the b60 μm particle size these sedimentological, mineralogical and geochemical fraction were used in this study. Samples were then kept studies were interpreted to record isotopically heavier in ∼40 mL of 0.5 N HCl solution over night to remove meteoric water δ18O values, temperatures increasing by any admixed carbonates and then rinsed with successive as much as 10 °C, and a landscape that became pro- aliquots of deionized H2O until the pH of the solutions gressively drier from Late Pennsylvanian to Early associated with the samples was equivalent to the initial Permian time. Significantly, these lithological and pH of deionized H2O. Each sample was subsequently geochemical palaeoenvironmental proxies appear to be treated over a period of 28–35 days with successive consistent with the biological record of changing fossil 40 mL aliquots of 30% H2O2 at room temperature to N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 163 remove organic matter. As the reactivity of the solution scanning mode from 2–70° 2Θ at 0.01° 2Θ/min, 40 kV diminished and the suspended particles settled, the and 44 mA. For Fe-oxide samples that contain goethite, solution was decanted and replaced by a fresh aliquot of the amount of Al substituted for Fe in goethite was H2O2. After H2O2 treatments, samples were dried in a determined by the XRD method of Schulze [28]. vacuum desicator at room temperature. In this paper, For chemical analysis, bulk samples were combined samples subjected only to the foregoing treatments are with lithium tetraborate to produce a 2:1 mixture on a designated bulk samples. mass basis. These mixtures were fused in graphite The mineralogy of the bulk samples was determined crucibles at temperatures of 1000 °C for one hour and by X-ray diffraction (XRD) analysis using Cu–Kα then quenched in deionized water to produce a glass that radiation on a Rigaku Ultima III X-ray diffractometer in was subsequently ground to b60 μm. Approximately the Department of Geological Sciences at Southern 125 mg of fused glass from each sample was then sealed Methodist University. Samples were back-mounted into in 15 mL Teflon bombs with 10 mL of concentrated an aluminum holder and analyzed under continuous HNO3 and left on a hot plate at 100 °C until all of the

Fig. 3. X-ray diffractograms of powdered permineralized plant material from (A) Sample F and (B) Sample C. Peak positions that were used for mineralogical identification are marked by dashed lines. Spacing (for Cu–Kα, given in angstroms) and the likely mineral responsible for the peak is given above the dashed lines. Peaks near 2.349 Å and 2.034 Å are from the aluminum holder used during analysis, and do not represent the contents of the samples. All of the identifiable peaks in Sample F are attributed to d(hkl) spacings of hematite, whereas Sample C includes peaks with d(hkl) spacings indicating the presence of goethite and phyllosilicate. See text for discussion. 164 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171

solids dissolved. Dissolution of all of the samples took determined by measurement of H2O evolved during place within 2 days. Each 10 mL aliquot was then high-temperature dehydration of the samples. Samples transferred to 100 mL volumetric flasks and the solution were initially outgassed at low temperature (∼120 °C) − was diluted to 2 to 3% HNO3 . Chemical analyses of the for ∼10 h in order to remove sorbed water. Samples dilute HNO3+ sample solutions were performed on an were then heated to ∼1100 °C under vacuum to ICP-OES at the DANR Analytical Facility at the evolve structural hydroxyl water. The liberated struc- University of California, Davis. The relative analytical tural water was then converted to H2 by passage over error of these analyses is b±2% of the reported value for hot (∼750 °C) U-metal [31]. After complete conver- the oxide component. sion of evolved H2OtoH2,H2 gas yields were In order to assess the isotopic composition of non- measured in a mercury manometer with an uncertainty iron oxide constituents within the samples, ∼200 mg of ±1 μmol. aliquots of bulk samples were subjected to Citrate– Bicarbonate–Dithionite (CBD) digestion solutions at 4. Results ∼25 °C in order to remove iron-oxides. The remaining non-iron oxide residue was then washed with 5 to 4.1. Mineralogy 8 successive 50 mL aliquots of deionized H2O and three to four successive 40 mL aliquots of 30% H2O2. After The mineralogical composition of the bulk samples, H2O2 treatment, the residues were dried in an oven at as deduced from XRD analysis, is presented in Table 1. 40 °C. In this paper, these non-iron oxide constituents All of the samples have prominent peaks near 3.004 Å, are designated residue samples. 2.71 Å, 2.53 Å, 2.21 Å, 2.06 Å, 1.85 Å, 1.70 Å and Oxygen isotope analyses of the bulk and residue 1.60 Å, corresponding to the d(hkl) indices of hematite. samples were performed using BrF5 reagent following In addition, sample C (Table 1) exhibits relatively minor the procedure of [29] to produce O2 gas at the peaks near 4.19 Å, 2.45 Å, 2.24 Å, and 1.56 Å, Department of Geological Sciences at the Southern corresponding to the d(hkl) indices of goethite (Fig. 3). Methodist University. The O2 gas was then quantita- Small peaks near 7.21 Å, 3.58 Å and 3.34 Å correspond tively converted to CO2 and analyzed for oxygen to the d(00l) and d(002) indices of kaolinite and the d isotope composition on a Finigan MAT 252 IR Mass (333) index of quartz, respectively. Spectrometer. Results are reported in per mil notation, where 4.2. Chemistry R d18O ¼ sample −1 X 1000 Results of the chemical analyses of the bulk samples Rstandard are reported in Table 2 as the mole fraction of the oxide components. There is some oxygen that is derived from R= 18O/16O. δ18O values are reported relative the structural hydroxyl groups associated with goethite that standard mean ocean water (SMOW) [30]. Replicate must be taken into account in order to calculate the end- analyses of NBS 28 (n=4) indicate analytical errors of member δ18O value for Fe-oxide in Sample C (Fig. 3; b0.2%. Table 1). The mole fraction of oxygen from H2O(X Wt.% H2O values of samples containing goethite (O)H2O) in sample C, as determined from high-tem- (as determined by X-ray diffraction analysis) were perature dehydration, is reported in Table 2. The wt.%

Table 2 Mole fraction of oxygen of the various oxide components in each sample and measured δ18O values of bulk and residue samples 18 18 Sample X(O)P X(O)K X(O)Ca X(O)Mg X(O)Na X(O)Mn X(O)Fe X(O)Al X(O)Si X(O)Ti X(O)H2O X(O)Hem δ Obulk δ Oresidue (%) (%) H 0.0220 0.0002 0.0088 0.0000 0.0004 0.0063 0.9423 0.0132 0.0068 0.0000 0.0000 0.9423 4.8 22.0 G 0.0032 0.0004 0.0001 0.0001 0.0001 0.0082 0.9780 0.0060 0.0034 0.0006 0.0000 0.9780 2.1 18.3 F 0.0136 0.0005 0.0005 0.0001 0.0001 0.0063 0.9673 0.0051 0.0062 0.0004 0.0000 0.9673 1.9 16.4 E 0.0001 0.0004 0.0001 0.0001 0.0001 0.0144 0.9780 0.0041 0.0025 0.0004 0.0000 0.9780 1.5 17.7 D 0.0070 0.0001 0.0007 0.0000 0.0000 0.0055 0.9566 0.0249 0.0051 0.0002 0.0000 0.9566 1.4 20.6 C 0.0040 0.0005 0.0001 0.0001 0.0001 0.0099 0.8891 0.0122 0.0116 0.0003 0.0722 0.9549 1.3 32.1 B 0.0044 0.0004 0.0001 0.0001 0.0001 0.0100 0.9697 0.0090 0.0062 0.0001 0.000 0.9697 0.4 17.6 A 0.0013 0.0009 0.0001 0.0001 0.0001 0.0106 0.9506 0.0101 0.0258 0.0005 0.000 0.9506 1.7 17.7 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 165

structural H2O in sample C was determined before (2.6) The mole fraction of oxygen from Fe-oxide minerals (X and after (5.8) Citrate–Dithionite treatment. This corre- (O)Hem) in all samples analyzed represents ∼94% to 98% sponds to 90.2% of the oxygen in structural H2O of the total oxygen derived from the sample during BrF5 originating from hydroxyl groups in goethite from sample reaction. C. The remaining 9.8% is likely derived from structural Measured δ18O values for the bulk and residue hydroxyl originating in kaolinite (Table 1, Fig. 3). samples are presented in Table 2. These δ18O values Therefore, in addition to oxygen derived from Fe–O range from 0.4% to 4.8%. In all cases, the residues bonds (X(O)Fe;i.e.onlyFe–O bonds in hematite and remaining after complete dissolution of Fe-oxides have goethite), sample C has an additional 0.065 mole fraction more positive δ18O values than the bulk (hematite-rich) of oxygen from structural hydroxyl in goethite. Thus, the fractions. This probably reflects the generally more mole fraction of total oxygen from all Fe-oxides in positive range of δ18O values of other common minerals sample C (X(O)Hem) is 0.9451. All of the other Fe-oxide (e.g., phyllosilicates and quartz) compared to Fe (III) samples are hematite and their X(O)Fe values are oxides in these samples [4,7].Inthisregard,the 18 equivalent to their X(O)Hem values, because all of the difference between the δ O of X(O)Fe values measured oxygen in this mineral occupies Fe–O bonds (i.e., Fe2O3). for the bulk (hematite-rich) and residue (CD-treated)

Table 3 Stratigraphic position and δ18O values measured for paleosol hematite and phyllosilicate samples Position Hematite 103ln18α= Hematite Hematite Phyllosilicate 103ln18α= Phyllo T (°C) Phyl-Hem T (°C) Phyllosilicate (m) a sample δ18O δ18O Oxygen isotope pair δD and δ18O (%) (%) 850 1.63⁎106/T 2 −12.3 b H 3.7 ––– 0.413⁎106/T 2 −2.56 c 0.773⁎106/T 2 −6.194 d 800 1.63⁎106/T 2 −12.3 b G 1.7 ––– 0.413⁎106/T 2 −2.56 c 0.773⁎106/T 2 −6.194 d 750 1.63⁎106/T 2 −12.3 b F 1.4 ––– 0.413⁎106/T 2 −2.56 c 0.773⁎106/T 2 −6.194 d 710 1.63⁎106/T 2 −12.3 b E 1.1 2.82⁎106/T 2 −5.06 22 25±1 27 0.413⁎106/T 2 −2.56 c 49±1 0.773⁎106/T 2 −6.194 d 54±1 620 ––2.83⁎106/T 2 −4.73 22.6 28 525 ––2.83⁎106/T 2 −4.78 22.7 29 475 ––2.85⁎106/T 2 −4.85 21.2 34 415 ––2.83⁎106/T 2 −5.04 21.8 26 395 ––2.82⁎106/T 2 −5.00 21.1 33 365 ––2.83⁎106/T 2 −5.13 21.1 34 315 1.63⁎106/T 2 −12.3 b D 0.5 2.84⁎106/T 2 −5.08 20.6 35 35 0.413⁎106/T 2 −2.56 c 56 0.773⁎106/T 2 −6.194 d 62 265 2.83⁎106/T 2 −4.87 20.8 31 230 1.63⁎106/T 2 −12.3 b C −0.4 2.83⁎106/T 2 −5.23 19.5 37 34 0.413⁎106/T 2 −2.56 c 56 0.773⁎106/T 2 −6.194 d 60 205 1.63⁎106/T 2 −12.3 b B −0.1 2.83⁎106/T 2 −6.20 19.6 26 33 0.413⁎106/T 2 −2.56 c 50 0.773⁎106/T 2 −6.194 d 54 180 1.63⁎106/T 2 −12.3 b A 0.9 2.76⁎106/T 2 −6.75 19.6 24 24 0.413⁎106/T 2 −2.56 c 50 0.773⁎106/T 2 −6.194 d 54 135 ––2.78⁎106/T 2 −6.11 20.4 – 22 a Stratigraphic height above the basal Markley Formation boundary. b Yapp [3]. Synthetic hematites, pH=1–2. c Clayton and Epstein [2]. Natural assemblage formed from alkaline solution. d Bao and Koch [1]. Synthetic hematites, pH=8–9. 166 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 sample fractions represents two points upon a two-end- time [34]. However, in order for these minerals to be member mixing line, where: utilized as palaeothermomters, the temperature-dependent oxygen isotope fractionation between the various miner- d18 ¼ ð Þ ⁎d18 þ ð Þ ⁎d18 als and coexisting water must be known. Obulk X O Hem OHem X O residue Oresidue 5.1. Mineral-water oxygen isotope fractionation factors and 1 ¼ XðOÞ þ XðOÞ Hem residue As discussed by Tabor and Montañez [18], the phyllosilicate-water oxygen isotope fractionation equa- These two end-member mixing relationships result in tions for the Permo-Pennsylvanian samples vary calculated oxygen isotopic compositions for the end- according to the mole fraction of oxygen contributed − member Fe-oxide minerals from 0.4% to 3.7% (Table 3; from the 1:1 and 1:2 phyllosilicate minerals present δ18 Fig. 4). Specifically, the most negative O values occur within the samples. Furthermore, the chemical compo- within the Late Pennsylvanian strata of the Eastern sition of the 2:1 phyllosilicates within each sample has Midland basin, whereas the Early Permian strata exhibit an effect upon phyllosilicate-water oxygen isotope δ18 heavier Ovalues(Fig. 4). Although portions of the fractionation. The calculated phyllosilicate-water oxy- δ18 Early Permian strata are not represented by hematite O gen isotope fractionation factors for the Permo-Penn- data, oxygen isotopic compositions among Fe-oxides, sylvanian phyllosilicates are presented in Table 3. phyllosilicates and calcites exhibit similar stratigraphic Hematite-water oxygen isotope fractionation equa- patterns through Upper Pennsylvanian and Lower tions have been presented by several authors at different Permian strata (Fig. 4).

5. Discussion

Similar intrabasinal stratigraphic trends have been documented for δ18O values of phyllosilicates and calcites [11,19], phyllosilicates and iron-oxides [16,32] as well as calcites and iron oxides [33]. These studies assert that similar stratigraphic δ18O trends would only be expected between two minerals if they independently record similar temporal changes in environmental factors such as δ18O values of regional waters (i.e., rainfall and/or groundwa- ter) and Earth-surface temperatures. As mentioned earlier, Permo-Pennsylvanian calcites and phyllosilicates from the eastern Midland basin exhibit similar stratigraphic trends indicative of a regional transition to isotopically heavier soil water from Pennsylvanian through Permian time. However, the paleopedogenic calcite and phyllosilicate data also exhibit oxygen isotope disequilibrium, because the calcite samples crystallized from evapora- tively modified soil waters that were enriched in O18 [18,19].Therefore,theΔ18O values between contempo- raneously-formed pedogenic calcite and phyllosilicate δ18O values are not permissive of oxygen isotope pair Fig. 4. Plot showing the stratigraphic position (in meters above the paleothermometry. Nevertheless, The Permo-Pennsylva- δ18 18 base of the Markley Fm.) versus the measured O values of hematites nian phyllosilicates and Fe-oxide δ O values also exhibit (filled circles; Table 3), Phyllosilicates (filled squares; [17,18]; Table 3) similar stratigraphic trends (Fig. 4), suggesting they and calcites [17,18] sampled from palaeosols in the Upper Pennsylvanian record similar environmental conditions at or near isotope and Lower Permian strata of north-central Texas. The horizontal dashed – equilibrium. If so, paired oxygen isotope values of the line corresponds to the approximate location of the Pennsylvanian Permian boundary in the study area. All three minerals show a general phyllosilicate and hematite minerals may provide stratigraphic trend toward heavier δ18O values from Pennsylvanian geothermometers to indicate temperatures of mineral through Permian strata, suggesting a change toward generally heavier formation at Earth's surface during Permo-Pennsylvanian δ18O values of soil moisture and rainfall through this time. See text. N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 167 ranges of temperatures and pH's for both natural and values results in a calculated temperature for equilibri- synthetic samples [1–3], Table 3: um coprecipitation:

1000 ln18a ¼ 0:413⁎106=T 2−2:56ðnatural assemblage½2Þ ð Þ 3 18a ¼ : ⁎ 6= 2 − : ½ 1 10 ln phyllo 2 82 10 TK 5 06 18

3 18a ¼ : ⁎ 6= 2 − : ½ 10 ln Hem 1 63 10 TK 12 3 3 1000 ln18a ¼ 0:542⁎106=T 2−5:221ðpH ¼ 12; ½1Þ ð2Þ

3 18a − 3 18a ¼ : ⁎ 6= 2 − : 10 ln phyllo 10 ln Hem 2 82 10 TK 5 06 − : ⁎ 6= 2 − : 1 63 10 TK 12 3 18 ⁎ 6 2 3 18 18 1000 ln a ¼ 0:733 10 =T −6:194ðpH ¼ 89; ½1Þ ð3Þ Let; 10 lnð1000 þðd Ophyllo−d OHemÞ=1000Þ 3 18 3 18 ¼ 10 ln aphyllo−10 ln aHem ð5aÞ

18a ¼ : ⁎ 6= 2− : ð ¼ ½ Þ ð Þ 3 ð : Þ¼ð : − : Þ⁎ 6= 2 þð− : þ : Þ ð Þ 1000 ln 1 63 10 T 12 3 pH 1 2; 3 4 10 ln 1 0209 2 82 1 63 10 TK 5 06 12 3 5b

These studies indicate significantly different hema- 20:68 ¼ 1:19⁎106=T 2 þð7:24Þð5cÞ tite-water oxygen isotope fractionation at low tempera- K tures (Fig. 6). In particular, it appears that hematite- water oxygen isotope fractionation factors may be 1:19⁎106=13:44 ¼ T 2 ð5dÞ sensitive to solution pH [1,35]. There are very few K data for hematite and coexisting water δ18O values from young, natural environments where temperatures and M 2 ¼ M ð Þ TK 88541 5e pH levels are known [35]. However, Girard et al. [16] published hematite and water δ18O values from lateritic soil profiles in Yaou, French Guinea, where soil pH's are TK ¼ 298 ð5fÞ relatively acid (b4.5) and soil mean annual temperatures are ∼25 °C. Oxygen isotope fractionation between ð- Þ¼ F : ð Þ those hematites and local meteoric water exhibits a large T C 25 3 5g range of 103lnα values from 5.8% to 11.0% at 25 °C. (Fig. 6). However, 91% of the Yaou samples exhibit less This approach was used to calculate oxygen isotope variation, with hematite-water 103lnα values ranging pair estimates of palaeotemperature between δ18O from 6.8±0.8% (1 s). Such 103lnα values compare most values of penecontemporaneous hematite and phyllosi- favorably with the hematite-water oxygen isotope licate samples. The reported uncertainty of the calculat- fractionation value of 6.0% at 25 °C proposed in Yapp ed temperature of crystallization reflects the sum of ([35]; 1990; Eq. (4); Fig. 5). Based upon this example errors associated with the fractionation equations, the (and several others discussed in [35]), the hematite- mole fractions of the oxide components and the δ18O water oxygen isotope fractionation equation of Yapp [3] measurement. The results of these calculations, includ- is adopted here. ing different hematite-water oxygen-isotope-fraction- As mentioned, if Permo-Pennsylvanian hematites ation equations discussed above, are reported in Table 3. and phyllosilicates from north central Texas record For the Yapp [3] Hematite-water oxygen isotope conditions of oxygen isotope equilibrium, then the fractionation equation, hematite-phyllosilicate oxygen isotopic difference between penecontemporaneously isotope pairs appear to be consistent with isotope equi- formed hematites and phyllosilicates may provide an librium at Earth-surface conditions, with temperatures estimate of temperature at the time of mineral of co-precipitation ranging from 24 to 35 °C (Table 3). precipitation (Fig. 6). For example, at a stratigraphic When the Clayton and Epstein [2] and Bao and Koch [1] level of ∼710 m (Table 3), phyllosilicate and hematite hematite-water oxygen isotope fractionation equations samples have δ18O values of 22.0±0.2% and 1.1± are considered for the same penecontemporaneous 0.2%, respectively. Combination of the hematite (Eq. (4) phyllosilicate-hematite oxygen isotope pairs, apparent above; [3]) and phyllosilicate (Table 3; [18]) oxygen equilibrium temperatures are too high for Earth-surface isotope fractionation factors with the measured δ18O conditions, ranging from 50–62 °C. 168 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171

It is difficult, probably even impossible, to assess whether ancient minerals truly record isotope equilibrium, or whether they reflect different environments of formation that coincidentally provide reasonable palaeotemperature estimates. As mentioned, Tabor and Montañez [18] estimated the Permo-Pennsylvanian palaeotemperatures from palaeosol profiles across the Permo-Pennsyl- vanian landscape using the combined δ18OandδDvalues of Permo-Pennsylvanian palaeopedogenic phyllosilicates. These palaeotemperature estimates are given in Table 3. There are five stratigraphic horizons where hematite-phyllosilicate oxygen isotope pairs and phyllo- δ18 δ silicate Oand D values are available for independent Fig. 6. Plot of the oxygen isotope fractionation values (α-values) for estimates of palaeotemperature (Table 3). The hematite- Hematite and Phyllosilicate vs. temperature in degrees Kelvin. α values for phyllosilicate oxygen-isotope pair temperature estimates hematite follow the oxygen isotope fractionation equation of Yapp (1990), based on the hematite-water oxygen isotope fractionation whereas α values for phyllosilicate follow the oxygen isotope fractionation equation for smectite with the chemical composition collected from 710 m equations of Clayton and Epstein [2] and Bao and Koch Δ ∼ – (Table 3; [18]). The values measured between oxygen isotope values of [1] are consistently 20 25 °C higher than stratigraphi- co-precipitated hematite and phyllosilicate correspond to equilibrium cally equivalent palaeotemperature estimates based on temperature of crystallization. The Δ value between paleopedogenic phyllosilicate δ18OandδD values. However, when the hematite and phyllosilicate from 710 m is 0.0210, which corresponds to an hematite-water fractionation equation of Yapp [3] is used equilibrium temperature of crystallization of 25±3 °C. See text for details.

to estimate palaeotemperatures, all five horizons have phyllosilicate-hematite oxygen isotope pairs and phyllosilicate δ18OandδD palaeotemperature estimates that are analytically indistinguishable from one another (Fig. 7, Table 3).

5.2. Palaeoenvironmental implications of Permo- Carboniferous temperature change

Unless these results are coincidental, the similar estimates for temperatures of crystallization that are provided from δ18OandδD of the Permo-Carboniferous palaeopedogenic phyllosilicates [18] to those of oxygen isotope pairs between coexisting palaeopedogenic hematite and phyllosilicate (using equation of [3]; Fig. 7) have several important implications: (1) hematite and phyllosilicate samples taken from the palaeosol profiles likely preserve a record of oxygen isotope equilibrium, (2) the single- mineral palaeotemperature estimates based upon δ18O and δD of phyllosilicate values assume equilibrium conditions with waters that lie upon the meteoric water line of Craig ([36]; δD=8⁎δ18O+10). Given the agreement 3 18α between the single-mineral and oxygen-isotope pair Fig. 5. Plot of oxygen isotope fractionation (10 ln ) versus temperature 18 (as 106/T2) values with calculated hematite-water oxygen isotope estimates for palaeotemperature, both hematite δ Oand 18 fractionation curves from [1–3]. Open squares represent 103ln18α values phyllosilicate δ OandδD likely preserve isotope values for natural soil hematites from 3 different units collected in Yaou that represent equilibrium with meteoric water. This δ18 ‰ Province, French Guinea [16].Themodernmeteoric Ovalue,3.0 , strongly suggests that controls upon the isotope compo- was used to calculate 103ln18α values for Unit 1. Meteoric water δ18O values used for hematite samples in units 2 and 3 are −4.0% and −5.0%, sition of Permo-Pennsylvanian meteoric precipitation (i.e, respectively, which are the interpreted meteoric water δ18Ovaluesatthe oceanic isotope composition, kinetic and equilibrium time unit 2 and unit 3 hematites formed [16]. fractionation in meteoric precipitation) are similar to N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 169

tures can only be confidently discussed as a temperature rise that is minimally 5 °C (from 27 °C to 32 °C) from Pennsylvanian to Permian time. Coincidentally, a rise of 5 °C is thought to be the range of Pleistocene and Holocene temperature change in the tropical latitudes between glacial and interglacial states (although the modern tropics never got quite so hot as those reported for the Permian) [42]. Finally, although some of the paleotemperature estimates from Permian strata may seem too high to support an ecosystem and sustain plant life, it has been known for nearly a century [e.g., 43,44] that the temperature at which photosynthesis is most rapid ranges between 29 and 32 °C, and there is an exponential decline in plant growth away from that optimum range of temperatures. Within analytical uncertainty, the most warm paleotemperature estimates overlap with this range of optimum temperatures for photosynthesis. Therefore, the real limitation to plant growth and paleoecosystem health at these estimated paleteomperatures was likely the availability of water for photosynthesis, rather than the temperature itself. Upon consideration of these issues, the reported range of Permo- Pennsylvanian paleotemperature estimates is considered Fig. 7. Plot of the calculated temperatures based on (1) hematite– reasonable, and they are within the range of permissible phyllosilicate oxygen isotope pairs (large gray filled circles) and values for modern tropical ecosystems. 18 (2) phyllosilicate δ O and δD values assuming equilibrium with Significant palaeotemperature change has been pre- meteoric waters (small open circles) relative to their stratigraphic dicted from Permo-Carboniferous General Circulation position (in meters) above the base of the Late Pennsylvanian Markley Fm. The horizontal dashed line corresponds to the approximate Models if atmospheric PCO2 increased from near present location of the Pennsylvanian–Permian boundary in the study area. atmospheric levels (PAL) during Carboniferous time to 4X See text for discussion. to 8X PAL during Permian time [38], and references therein. Significantly, some evidence does exist for in- modern controls upon the meteoric hydrological cycle creasing atmospheric PCO2 change, from ∼1to10XPAL, [36,37]. (3) The isotope composition of the studied from Carboniferous to Permian time [39,40], which ap- hematites and phyllosilicates retain their original Permo- pears to be temporally and mechanistically linked with the Pennsylvanian values. (4) The significant temperature Late Palaeozoic Gondwanan deglaciation [40]. change (∼6–10 °C) through Permo-Carboniferous time If such a temperature change occurred, it would have that is indicated from these data may be taken seriously. likely had a profound impact upon the ecology of this Nevertheless, some of the Permian paleotemperature region. For example, elevated atmospheric temperature estimates presented here may seem implausible, because results in greater potential evapotranspiration and loss of they are significantly higher than modern tropical tem- plant-accessible soil moisture [45,46]. Considering the peratures, and such conditions may be incapable of general observation that potential evapotranspiration supporting plant ecosystems. However, it is noteworthy increases by 0.2 mm/d for every 1 °C increase in that soil subsurface horizons are buffered against seasonal temperature [46],a6–10 °C temperature change would swings in temperature, and mean annual subsurface soil correspond to an increase in potential evapotranspiration temperatures are typically 1° to 2 °C warmer than mean of ∼400–700 mm/yr. Therefore, in the absence of a annual surface air temperature [41]. The paleotemperature concomitant increase in rainfall to offset increased estimates provided here are interpreted to represent the soil evapotranspiration, soil moisture content and soil water subsurface and, as such, should be regarded as maximum, storage would have become significantly drier and less, or slight overestimates, of mean annual surface atmo- respectively, from Pennsylvanian to Permian time. spheric temperatures. In addition, soil paleotemperature Furthermore, this palaeotemperature change would estimates from the oxygen isotope pair analyses range have resulted in less surplus water, or through-flow in from24±3°Cto35±3°C.Duetotheanalyticalun- soil profiles, thus reducing (1) leaching and chemical certainty associated with these estimates, paleotempera- weathering in the soil profile and (2) recharge to local 170 N.J. Tabor / Earth and Planetary Science Letters 253 (2007) 159–171 and regional water tables. Significantly, the stratigraphic [7] C.J. Yapp, Climatic implications of surface domains in arrays of interval over which the estimated Permo-Pennsylvanian dD and d18O from hydroxyl minerals: Goethite as an example, Geochim. Cosmochim. Acta 64 (2000) 2009–2025. temperature change occurred preserves the stratigraphic [8] C.J. Yapp, H. Poths, 13C/12C ratios of the Fe(III) carbonate transition from poorly drained swampland deposits with component in natural goethites, Geochem. Soc. Spec. 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