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Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 Contents lists available at ScienceDirect Palaeogeography, Palaeoclimatology, Palaeoecology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o Carbon and oxygen isotope evidence for high-frequency (104–105 yr) and My-scale glacio-eustasy in Middle Pennsylvanian cyclic carbonates (Gray Mesa Formation), central New Mexico Maya Elrick ⁎, Lea Anne Scott Earth and Planetary Sciences, University of New Mexico, Albuquerque NM 87131, United States a r t i c l e i n f o Article history: Received 16 July 2009 Received in revised form 5 November 2009 Accepted 17 November 2009 Available online 22 November 2009 Keywords: Pennsylvanian Carbon and oxygen isotopes Glacio-eustasy Conodonts Icehouse climates a b s t r a c t We combine cyclo- and sequence stratigraphy along with whole rock δ13C and conodont apatite δ18O analysis to document high-frequency (104–105 yr) and My-scale sea-level changes for the Middle Pennsylvanian (Desmoinesian or Moscovian) Gray Mesa Formation of central New Mexico. Approximately 75 subtidal cycles (1–8 m) are grouped into 4 1/2 My-scale depositional sequences (40–80 m). About 50% of the cycles show evidence of prolonged subaerial exposure at cycle tops with the development of calcretes, diagenetic mottling, and regolith intraclasts. High-resolution δ13C analysis of whole rock limestones across nine of the cycles indicates that the cycle tops were diagenetically altered by isotopically light, meteoric fluids during sea-level fall and lowstand. These δ13C trends support the interpretation that high-frequency sea-level changes were responsible for cycle development. Conodont apatite δ18O values from sampled cycles indicate that the high-frequency sea-level changes were driven by glacio-eustasy combined with changes in surface seawater temperature (SST). δ18O values from conodont apatite, spanning parts of three depositional sequences indicate that My-scale glacio-eustasy and/or SST changes controlled sequence development. δ18O shifts indicate that the magnitudes of 104–105 yr glacioeustasy were between ∼ 55 and 170+ m combined with tropical SST changes of ∼ 1.5°–6 °C. Calculated My-scale glacio-eustatic oscillations were between ∼ 60 and 140 m with SST changes of b 3.5 °C. The most plausible driver for the My-scale paleoclimate changes is long-period obliquity (∼ 1.2 My) variations. These calculated high-frequency, glacio-eustatic values are similar or greater than Pleistocene values, and lie within the range estimated for other Middle Pennsylvanian successions using a variety of independent eustatic proxies. The similarity in range of magnitudes between high-frequency and My-scale sea-level changes combined with the large differences in magnitudes between individual high-frequency sea-level oscillations helps explain the lack of systematic cycle-stacking patterns within these Pennsylvanian icehouse sequences. © 2009 Elsevier B.V. All rights reserved. 1. Introduction The Late Paleozoic (Carboniferous to Permian) ice age is well documented by the occurrence of glacial deposits on all continental remnants of southern Gondwana (Caputo and Crowell, 1985; Veevers and Powell, 1987; Frakes and Francis, 1988; Crowell, 1999; Isbell et al., 2003; Fielding et al., 2008). Traditional reconstructions of this glacial period invoke a single, large, and long-lived (∼50–90 My) ice sheet (e.g., Veevers and Powell, 1987). More recently, the concept of discrete glacial and nonglacial intervals each lasting on the order of several million years and involving multiple ice sheets is unfolding (Isbell et al., 2003; Montañez et al., 2007; Fielding et al., 2008). Far-field (low-latitude) evidence of the waxing and waning of these ice sheets comes from the globally widespread occurrence of ⁎ Corresponding author. E-mail addresses: [email protected] (M. Elrick), [email protected] (L.A. Scott). 0031-0182/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2009.11.023 high-frequency (104–105 yr) cycles/parasequences (or cyclothems). In the Pennsylvanian of the U.S. Midcontinent and United Kingdom, these high-frequency cycles are typically composed of a combination of offshore siliciclastics, nearshore skeletal limestones, fluvial-deltaic siliciclastics, coals, and paleosols (e.g., Moore, 1964; Ramsbottom, 1979; Maynard and Leeder, 1992; Heckel, 1977, 1994; Wright and Vanstone, 2001). In contrast, many Pennsylvanian marine successions in the U.S. Southwest are composed predominantly of offshore through nearshore carbonates; nonmarine siliciclastics, paleosols, subaerial exposure features can be either poorly exposed or relatively weakly developed (Goldhammer and Elmore, 1984; Algeo et al., 1991; Soreghan, 1994; Wiberg and Smith, 1994; Scott and Elrick, 2004). The magnitude of sea-level changes associated with the development of these Pennsylvanian cycles has been estimated using a variety of independent mechanisms including ice-volume modeling (Crowley and Baum, 1991), facies juxtaposition (Heckel, 1994, 1997), paleotopography (Soreghan and Giles, 1999), and δ18O isotopic shifts (Adlis 308 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 et al., 1988; Joachimski et al., 2006). In this paper, we 1) present highresolution δ13C trends associated within Middle Pennsylvanian cycles and cycle-capping early diagenetic features to verify high-frequency 18 sea-level changes, 2) present conodont apatite δ O trends from cycles and My-scale (3rd-order) depositional sequences to document glacioeustasy as the driver of 104–105 yr and My-scale sea-level changes, and 3) estimate the magnitudes of the glacio-eustatic changes and discuss the implications these magnitudes have for cycle-stacking patterns in icehouse climate modes. 2. Geologic setting and stratigraphy The early assemblage of Pangea in the Pennsylvanian resulted in the development of the Ancestral Rocky Mountains and associated intracratonic basins across the western U.S. (Kluth and Coney, 1981; Ye et al., 1996). During the Pennsylvanian, New Mexico lay within 5–10° of the paleoequator and was covered by shallow tropical seas interrupted by uplifted mountain blocks composed predominately of Precambrian crystalline rocks (Fig. 1a). Fig. 1. (a) Paleogeographic map of New Mexico during the Middle Pennsylvanian; inset of the western U.S. shows the location of New Mexico. Darker gray areas represent uplifts related to formation of the Ancestral Rocky Mountains. Mesa Sarca study area in black box. Other Middle Pennsylvanian outcrops discussed in text include: MA = Mesa Aparejo, SM = Sandia Mountains, SA = Sacramento Mountains. (b) Chronostratigraphic diagram for study area and most of central New Mexico. Vertical lines represent a hiatus. Gray Mesa Formation shown in shaded area. Biostratigraphic ages and nomenclature compiled from Kelley and Wood (1946) and Kues (2001). M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 The Lucero basin study area lay east of the low-relief Zuni uplift and accumulated up to ∼ 300 m of Middle Pennsylvanian carbonates and subordinate siliciclastics (Figs. 1a and 2). Northward thickening and coarsening of the siliciclastics within this basin suggest that most of them were sourced from the Penasco uplift lying north of Albuquerque (Fig. 1a). Based on extrapolation of data presented by Dickinson and Lawton (2003), subsidence rates in the Lucero basin were highest during the Late Pennsylvanian (Virgilian or Kasimovian) with tectonic movement occurring along high-angle normal faults (Cather, 2001). Within the study area, Pennsylvanian deposits lie unconformably above Precambrian crystalline rocks and are composed of coarse- to fine-grained transgressive siliciclastics (Sandia Formation), mixed carbonate-siliciclastics (Madera Group), and are conformably overlain by Permian fluvial red beds of the Bursum and Abo Formations 309 (Fig. 1b). The Middle Pennsylvanian (Desmoinesian or Moscovian) Gray Mesa Formation (∼ 290 m) is exceptionally well exposed at Mesa Sarca in central New Mexico (Figs. 2 and 3). At this locality, individual beds can be physically traced along depositional strike for over 4 km. Age control on the Gray Mesa Formation is based on limited fusulinid and preliminary conodont biostratigraphy (Kelley and Wood, 1946; Wengerd, 1959; Martin, 1971, this study). 3. Cycles and sequences 3.1. Meter-scale cycles Shallow subtidal through deep subtidal depositional facies are recognized within the Gray Mesa Formation; detailed descriptions and depositional interpretations of the various facies are reported in Fig. 2. Generalized stratigraphic column of Gray Mesa Formation showing depositional sequences, boundaries of high-frequency cycles, and sea-level curve interpreted from sequence stratigraphy. The sequence boundaries appear more abrupt that observed in the field because of scale of figure. Note the lack of thickening and thinning cycle-stacking 13 patterns and lack of systematic distribution of cycle-capping subaerial exposure features. See text for additional discussions. Filled triangles = cycles sampled for δ C analysis, open 18 triangles = cycles sampled for δ O conodont apatite analysis. 310 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 Fig. 3. Field photographs of Gray Mesa Formation at Mesa Sarca study site. (a) Depositional sequences 1–5 (white vertical lines) defined by slope-forming transgressive system tracts (TST) and maximum flooding zones (MFZ) and cliff-forming highstand system tracts (HST). Entire outcrop is about 350 m thick. (b) Typical high-frequency cycle (#12) with thin transgressive base (recessive under hang in shadow), thick-bedded shallow subtidal cap, and microkarst erosional top. Isotopic trends for this cycle shown in Figs. 5c and 7a. (c) Field photograph of parts of Sequences 2 and 3 (thick white lines) and tops of internal high-frequency cycles (thin white lines). Scott (2004) and Scott and Elrick (2004). The subtidal facies are arranged into ∼ 75 m-scale (average thickness of ∼3 m) upwardshallowing cycles and form subtidal cycles (in contrast to peritidal cycles which are capped by tidal flat or beach/foreshore facies). Deeper subtidal cycles are characterized by poorly exposed calcareous mudstone (deep subtidal facies) at the base overlain by skeletal wackestones through packstones/rare grainstones (shallow subtidal facies), whereas shallow subtidal cycles lack the calcareous mudstone at their base and are composed entirely of skeletal wackestones through packstones (Figs. 3–5). The cycles are asymmetric and record Fig. 4. Field photographs of cycle-capping early diagenetic features. (a) Black laminated calcite or pedogenic calcrete (white arrows) infilling horizontal cracks within cycle cap. (b) Reddish gray diagenetic mottles (white arrows) embedded in light gray, fine limestone matrix. (c) Subrounded to subangular intraclasts (white arrows) embedded in light gray, fine limestone matrix. M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 311 312 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 either entirely regressive deposition above the flooding surface or a thin transgressive unit followed by a thick interval of regressive deposits. Maximum flooding occurs near the base of cycles, though often within poorly exposed calcareous mudstone intervals, so it is not possible to observe the actual turnaround in depositional patterns. Approximately 50% of the cycle tops display early diagenetic features indicative of subaerial exposure (discussed below). The average cycle duration is estimated by dividing the length of the Desmoinesian/Moscovian by the number of observed cycles. Depending on which time scale is used (Ross and Ross, 1988; Klein, 1990; Heckel, 2003; Gradstein et al., 2004), the average cycle duration 13 ranges from ∼50 to 100 ky (“high-frequency”). This duration is an approximation because an unknown number of cycles occur within the underlying, poorly exposed earliest Desmoinesian Sandia Formation, and the exact position of the Middle–Upper Pennsylvanian boundary based on fusulinid biostratigraphy is poorly constrained in the study area (Martin, 1971). 3.2. Early diagenetic features Early diagenetic features at cycle caps and some cycle bases are subtle, but pervasive in outcrop (Fig. 4) and include: 1) dark gray to 13 Fig. 5. δ C trends from fine-grained matrix of high-frequency cycles. Refer to Fig. 2 for stratigraphic location of sampled cycles. (a) Cycles showing distinct up-cycle decreases in δ C 13 13 values. Some cycles show the lowest δ C values tens of cm below the cycle top. Note that the immediately overlying δ C value at base of cycle #39 shifts back to typical 13 Pennsylvanian marine values. (b) Cycles with no distinctive up-cycle trends in δ C values. Cycle #58 displays field evidence of subaerial exposure with diagenetic mottling and 13 intraclasts, and the δ C values of these cycle-capping diagenetic features are highly depleted relative to the immediately associated fine-grained matrix (see Fig. 6 and text for additional discussion). The top of cycle #57 shows evidence of depletion throughout the sampled cycle top. (c) Three cycles, each sampled at two localities separated by ∼200 m, 13 show up-cycle variations in δ C trends and magnitudes indicating along-strike variations in early diagenesis. Cycles #17 and #13 show significantly different values along strike with up-cycle depletions at one locality and not at the other. Tops of cycles #12 and #13 have black calcite, diagenetic mottling, intraclasts, or microkarstic erosion, but record up-cycle 13 13 δ C depletion at only one locality. Note that all δ C values for the individual cycle-capping diagenetic features show depletion relative to the adjacent fine-grained host limestone. M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 313 Fig. 5 (continued). black laminated calcite infilling vertical and horizontal cracks and encrusting bed tops, 2) red- to orange-gray mottles within a gray to yellow limestone matrix, and 3) subrounded to subangular limestone intraclasts floating in a fine limestone matrix (Scott and Elrick, 2004). Dark laminated calcite and mottling overprints primary depositional textures and the intensity of their development decreases downwards from the cycle tops, typically affecting only the upper few tens of cm of the cycle. The intensity of diagenetic overprinting varies along strike within a single layer, but the overprinting effects can be traced at least 2 km along depositional strike. These field relationships indicate that the cycle-capping features developed during early diagenesis before deposition of the overlying beds and are related to subaerial exposure. Half of the cycles (50%) display at least one cyclecapping early diagenetic feature, while nearly 20% of the cycles display two or more features. 3.2.1. Black laminated calcite Dark gray to black laminated calcite is observed in b20% of cycle tops and occurs as discontinuous bands (mm to cm thick and up to 20 cm long), wispy stringers, patches, and bed–top encrustations (Fig. 4a). The bands infill horizontal to vertical cracks, and the stringers create an anastomosing network imparting a brecciated appearance to the host limestone. Thick chert bands and cement-filled fenestrae are also associated with the dark calcite. In thin section, the calcite is characterized by dark brown laminated micrite which coats grains, lines cavities, and is associated with alveolar–sepal structures, dilation cracks, and rhizoliths. 3.2.2. Diagenetic mottling Approximately 25% of the cycles are capped by diagenetic mottling, which is characterized by rounded irregular patches of 314 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 reddish, orange, to dark gray limestone (mottles) surrounded by light gray to yellow-gray limestone host rock (Fig. 4b). Mottles are 0.25– 10 cm in diameter and are circular to elongate in cross section; they are similar in size to burrows observed throughout the Gray Mesa Formation but differ by their discoloration. The red (most common) to orange discoloration is most intense along the periphery of individual mottles, overall mottle discoloration decreases downward from cycle tops, and the intensity of discoloration varies along depositional strike. In thin section, mottles and host limestone boundaries are diffuse with mottles composed of slightly more interparticle cement and micron-size hematite crystals. 3.2.3. Intraclasts Intraclasts occur at the tops of b20% of the cycles and are characterized by subangular to subrounded dark gray limestone clasts embedded in a light gray to yellow-gray limestone matrix (Fig. 4c); both clast-supported and matrix-supported textures are observed. Clasts are generally poorly sorted ranging in size from b1 cm to 14 cm, are composed of fine-grained limestone, and have clearly defined edges with the associated host limestone. The clasts may be silicified or reddened, with more intense reddening occurring around clast edges. Intraclast beds display sharp irregular to planar contacts with underlying limestones or can have gradational contacts. Overprinting by diagenetic mottling sometimes makes it difficult to distinguish between diagentic mottles and discolored rounded intraclasts. In thin section, clast boundaries are irregular and sharp. 3.3. Depositional sequences The Gray Mesa cycles are bundled into 4 1/2 depositional sequences which are ∼ 40–80 m thick (Figs. 2 and 3). Transgressive system tract (TST) through maximum flooding zones (MFZ) are characterized by slope-forming intervals of deep subtidal cycles, whereas highstand system tracts (HST) are composed predominatly of cliff-forming, carbonate-dominated shallow subtidal cycles (Figs. 2 and 3; Scott and Elrick, 2004). No single stratigraphic horizon representing the sequence boundary is detected, rather a sequence boundary zone (SBZ) or zone of accommodation minimum is recognized by the turnaround from carbonate-dominated shallow subtidal cycles to siliciclastic-based, deep subtidal cycles. Sequence durations range from ∼0.7 to 1.4 My (3rd-order) depending on which Middle Pennsylvanian time scale is used (Ross and Ross, 1988; Klein, 1990; Heckel, 2003; Gradstein et al., 2004). Fig. 2 illustrates the cycle-stacking patterns within depositional sequences. Of particular interest is that unlike typical greenhouse climate cycle-stacking patterns (e.g., Goldhammer et al., 1993), the Gray Mesa cycles do not display systematic thickening and thinning patterns defining My-scale accommodation trends, nor are cycles displaying cycle-capping exposure features concentrated near the accommodation minima. Instead, the exposure features occur throughout the sequences and some of the thinnest cycles occur within TST and MFZ intervals. The distinctive slope- versus cliff-forming outcrop patterns defining sequences can be traced across the length of Mesa Sarca (N4 km) and into the next outcrop belt approximately 15 km to the north (Mesa Aparejo; Fig. 1a). In addition, coeval marine carbonates in the Sandia Mountains N100 km to the northeast and Sacramento Mountains ∼250 km to the south (Fig. 1a) are comprised of four to five Middle Pennsylvanian depositional sequences (Algeo et al., 1991; Wiberg and Smith, 1994; Krainer and Lucas, 2004). The similarity in cyclic facies types, sequence number, and sequence thickness suggests that Middle Pennsylvanian 3rd-order accommodation space trends are regional in scale, rather than due to local variations in tectonically driven subsidence. Ongoing sequence stratigraphic and biostratigraphic studies will aid in determining the specific regional extent of sequence correlations. 4. Methods 13 Samples for δ C analysis were collected from subtidal cycles and cycle-capping diagenetic features to assess whether the subtidal marine limestones were altered by meteoric fluids during sea-level fall and lowstand. Nine of the 75 cycles were sampled; two of the sampled cycles lack cycle-capping diagenetic features and were sampled to evaluate whether subaerial exposure could be detected 13 from δ C isotopic trends alone. In addition, three of the cycles were sampled at sites separated by N200 m to assess lateral variations in 13 δ C trends. Cycle-capping black laminated calcite, diagenetic mottles, and intraclasts along with their immediately associated host limestone were sampled at 16 different stratigraphic horizons to evaluate the cm-scale effects of subaerial exposure and diagenesis. 13 A total of 275 samples were analyzed for δ C analysis. Samples were collected every 30–50 cm in the lower parts of cycles and every 5–20 cm in the upper parts. Fine-grained matrix (composed dominantly of pellets, microspar, and micritic/microsparitic cement) was collected from thick section billets with a Dremel tool and 1 mm-wide drill bit to avoid obvious skeletal fragments, coarse cements, and veins and to obtain the necessary resolution between mottles, intraclasts, and host limestone. Oxygen isotopes from conodont apatite were sampled across two cycles and across parts of three depositional sequences to evaluate the origins of high-frequency and My-scale sea-level changes. Sequencescale samples were collected within the deepest water facies of ten successive cycles to minimize the effects of high-frequency isotopic variation related to cycle development. For stable isotopes of carbonate samples, CO2 gas was extracted by reaction with 100% phosphoric acid on a Thermoquest-Finnigan Gas Bench II automated preparation device and was measured using a Finnigan Delta Plus XL continuous flow isotope ratio mass spectrometer. Analyses are reported relative to VPDB (Vienna Peedee belemnite) and were routinely analyzed and compared to NBS 19 and Carrara Marble 13 standards. Average standard deviations on standards are ±0.1‰ for δ C 18 values and 0.2‰ for δ O values and reported relative to PDB. Samples for conodonts were processed using standard conodont concentration techniques (Sweet and Harris, 1989). Once the conodonts were concentrated, the apatite was converted to Ag3PO4 using a modified version of O'Neil et al. (1994) to ensure that only phosphorous-bound oxygen was analyzed. The Ag3PO4 crystals were analyzed in a high temperature TC-EA reduction furnace at 1450 °C in a He stream (e.g., Elrick et al., 2009). Isotope ratios of resultant CO are measured in continuous flow isotope ratio mass spectrometry using a Finnigan Mat 252 mass spectrometer. The precision of analyses based on long-term measurement of standards is b0.3‰ and is monitored by multiple analyses of several phosphate standards interspersed with samples. Analyses are reported relative to SMOW. 5. Results 18 5.1. Whole rock δ Ocarb 18 Whole rock carbonate δ O values from the cycles range from −1.0 to −9.9‰ (average = − 6.0‰). The average values are significantly 18 lower than estimated δ O Pennsylvanian seawater values (−1.5 and −3.0‰; Brand, 1982; Grossman et al., 1991; Algeo et al., 1992; Mii et al., 1999) and show no systematic trends with upward-shallowing 13 facies patterns or with δ C trends. Therefore, we interpret that the 18 δ Ocarb values reflect the effects of diagenetic alteration during late burial. 13 5.2. Whole rock δ C Estimated Middle Pennsylvanian marine δ13C values are between ∼3‰ and 5‰ (Brand, 1989; Grossman et al., 1991; Algeo, 1996; Mii M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 et al., 1999). Whole rock carbonate δ13C values from the Gray Mesa subtidal cycles range from 4.1‰ to −4.6‰ with the lowest values occurring in the upper b1 m of cycle tops. Six of the nine sampled cycles show up-cycle δ13C depletions of between 2‰ to 4‰ (Fig. 5). In half of these, the smallest isotopic values occur a few tens of cm below the cycle top; in the other half, the smallest values lie at the cycle top. 13 δ C values from the immediately overlying cycle base shift back to typical Pennsylvanian marine values (Fig. 5a). The two sampled cycles which lack field identified, cycle-capping subaerial exposure features 13 (cycles #1 and 58) show no systematic trends in δ C values (Fig. 5b). Two of the three cycles sampled ∼200 m apart (cycles #13 and 17) 13 show significant along-strike differences in δ C isotopic trends (Fig. 5c). 13 Fig. 6 illustrates the δ C values of cycle-capping diagenetic features and their immediately adjacent host limestone. Values from black laminated calcite are consistently lower than their immediately adjacent limestone host by between 0.3‰ and 6‰. In half the 13 sampled diagenetic mottles, the mottles have δ C values up to 3‰ lower than the adjacent host limestone; the other half have values 13 similar to their host. Intraclasts consistently have δ C values between 0.4‰ and 5‰ lower than their associated matrix. Of interest is that 13 δ C values from diagenetic features from some cycle tops show depletion even when the whole rock values do not (cycle #58; Fig. 5b). 13 315 18 5.3. Conodont δ Oapatite 18 The δ Oapatite values of the two sampled cycles range from ∼19‰ to 22‰ and isotopic values increase by ∼1.0‰ to 3.9‰ from the base 18 to the top of the cycles (Fig. 7a). While δ Oapatite values show 18 systematic up-cycle increases, δ Ocarb values from corresponding whole rock limestone show no systematic variations (Fig. 7a) 18 supporting the interpretation that δ Oapatite values reflect primary 18 marine seawater, while δ Ocarb values reflect the effects of diagenesis. 18 δ Oapatite values spanning parts of Sequences 1, 2, and 3 range from ∼17‰ to 19.5‰ (Fig. 7b). The values decrease and reach a minimum within the TST/MFZ and early HST, increase and peak within the middle HST, systematically shift to lower values within the upper HST, and reach the lowest values in the TST/MFZ and early HST of the overlying Sequence 3. The magnitude of isotopic shift from the top of the underlying Sequence 1 to the early HST of Sequence 2 is ∼1‰ and from the HST of Sequence 2 to the HST of Sequence 3 is ∼2.4‰ (Fig. 7b). The magnitude of these cycle- and sequence-scale isotopic shifts are minimum values because the study area was located along the inner to middle shelf; therefore, sediments representing final stages of sea-level fall, lowstand, and the initial rise did not accumulate at this location. Given this, the full extent of the oxygen isotopic shifts is not recorded at this location. Fig. 6. δ C trends for cycle-capping early diagenetic features black laminated calcite (pedogenic calcrete), diagenetic mottles, and intraclasts. All sampled calcretes and intraclasts 13 13 have δ C values that are less than that of the immediately adjacent fine-grained host limestone. The δ C values for diagenetic mottles are significantly less than the adjacent host in half the sampled horizons illustrating the cm-scale heterogeneity of pore-fluid flow. See text for additional discussion. 316 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 subaerially exposed, but cycle top diagenesis occurred in a rockdominated pore-fluid system or that the meteoric cement abundance was low or variably distributed (Goldstein, 1991). 6. Discussion 13 6.1. Whole rock δ C isotopic trends 13 We interpret that the marine to near marine δ C values at cycle bases and the lighter isotopic values in the upper portion of cycles represent the effects of early diagenetic alteration by isotopically light meteoric fluids during high-frequency sea-level falls/lowstands and subaerial exposure (e.g., Allan and Matthews, 1982; Beeunas and Knauth, 1985; Algeo et al., 1991; Goldstein, 1991; Joachimiski, 1994; Algeo, 1996; Immenhauser et al., 2002; Railsback et al., 2003; Theiling et al., 2007). In this interpretation, during sea-level fall and lowstand, marine sediments were progressively exposed to vadose and phreatic zone meteoric waters enriched in isotopically light soil CO2, and 13 micritic cement with low δ C values precipitated from these fluids. It is likely that the entire cycle thickness (and more) was exposed to meteoric waters during this time, but the isotopic signature of resultant pore fluids in the lower parts of the cycles was dominated by 13 the dissolution of marine carbonates whose δ C values were significantly higher than meteoric fluids (i.e., pore fluids were rockdominated); therefore, any early diagenetic micritic cements at cycle bases record marine to near marine isotopic values. In contrast, the 13 δ C values of pore fluids in the upper portion of cycles were dominated by isotopically light meteoric fluids (water-dominated pore fluids); therefore, the early diagenetic micritic cements in cycle 13 13 tops record low δ C values. We suggest that the increase in δ C values in the upper few tens of cm of some of the cycle tops may be the result of micritic cements precipitated into pores spaces after exposure and 13 during the subsequent transgression. As a result, the whole rock δ C values at some cycle tops are greater than underlying samples, reflecting a mix between the meteoric values developed during subaerial exposure and more marine values developed during the 13 subsequent transgression. Similar peak depleted δ C values lying some depth below cycle tops have been reported for Pleistocene through Ordovician cyclic successions (Allan and Matthews, 1982; Joachimiski, 1994; Driese et al., 1994; Algeo, 1996; Immenhauser et al., 2002; Railsback et al., 2003). 13 Along-strike variations in up-cycle δ C trends (Fig. 5c) indicate lateral heterogeneities in 1) pore-fluid isotopic composition, flow paths and flow rates, 2) sediment porosity/permeability, 3) abundance, distribution, and isotopic composition of meteoric cement phases, 4) vegetation cover and abundance, and/or 5) local topography. Centimeter-scale diagenetic heterogeneities are also implied given the large differences in δ13C values between host limestone and immediately adjacent diagenetic mottles (Fig. 6). Similar along-strike 13 and cm-scale variations in δ C values are reported by Goldstein (1991) and Theiling et al. (2007). 13 Those cycles which record constant marine or near marine δ C values from top to bottom (cycles #1, 12, and 58) may indicate that those particular cycles 1) were not subaerially exposed, 2) the diagenetically altered portion of the cycle was eroded during exposure and/or subsequent transgression, 3) little primary porosity was available for precipitation of meteoric cements, and/or 4) early diagenesis occurred in a rock-dominated pore-fluid system, therefore 13 the micritic cements record marine to near marine δ C values. The fact that some cycles display field evidence of subaerial exposure, but 13 lack δ C evidence of meteoric diagenesis (cycle #12) and some cycles 13 show up-cycle δ C decreases at one locality but not ∼ 200 m away (cycle #17), suggests that the marine sediments were, in fact, 18 6.2. Cycle-capping diagenetic features 6.2.1. Black laminated calcites The laminated black calcites are interpreted as soil carbonates (laminar calcretes) precipitated during pedogenesis and root calcification (Harrison and Steinen, 1978; Goldstein, 1988; Wright, 1994). 13 This interpretation is supported by their consistently low δ C values (soil/plant-derived CO2), association with rhizoliths, alveolar–sepal structures, and dilation cracks, and their similarity to previously reported calcretes (e.g., Harris and Nuna, 1975; Harrison and Steinen, 1978; Goldstein, 1988; Tucker and Wright, 1994). 6.2.2. Diagenetic mottling We interpret that the diagenetic mottles represent backfilling of burrows by slightly more porous/permeable sediment, followed by early micritic cementation from meteoric fluids during sea-level fall and lowstand (Horbury and Quing, 2004). The abrupt difference in 13 δ C values between some of the mottles and adjacent host limestone (Fig. 6) reflects the variable primary porosity/permeability and flow paths of pore fluids. 6.2.3. Intraclasts We interpret that the isotopically depleted intraclasts at cycle tops represent soil regolith clasts developed during subaerial exposure and meteoric diagenesis (e.g., Goldstein, 1988; Tucker and Wright, 1994). 13 6.2.4. Summary of δ C trends 13 The combined results from δ C trends spanning individual cycles and trends from individudal cycle-capping diagenetic features clearly supports our interpretation that cycles developed in response to sealevel changes and that during sea-level fall, lowstand, and early rise, marine sediments were subaerially exposed and altered by meteoric fluids. 18 6.3. Conodont δ Oapatite trends Analyses of oxygen isotopes from marine calcite and apatite are a well-established paleoclimate tool for evaluating variations in seawater temperatures and isotopic changes related to the growth and melting of continental glaciers. In particular, Joachimski et al. (2006) analyzed conodont apatite from Upper Pennsylvanian cyclothems/cycles in the U.S. Midcontinent and reported systematic 18 up-cycle positive δ O isotopic shifts associated with upwardshallowing facies trends. They interpret that up to 1.7‰ isotopic shifts are the result of the glacial ice-volume effect combined with cooling surface seawater temperatures. Using the Pleistocene as an icehouse analog, they interpreted Late Pennsylvanian interglacial– glacial sea-level changes of over 120 m combined with subtropical surface water temperature changes of ∼2–4 °C. 18 We interpret that the δ Oapatite isotopic shifts associated with Gray Mesa cycle and sequence development are also the result of glacioeustasy (ice-volume effect) and tropical seawater temperature changes. Isotopic changes due to the effects of evaporation and/or freshwater influx (“salinity effect”) are ruled out because the amount of evaporation/freshwater influx required to explain the measured 18 Fig. 7. (a) δ Oapatite (from conodont apatite) trends for sampled high-frequency cycles. Invariant δ O values from whole rock carbonate are also shown for comparison in cycle #12, 18 18 which supports the interpreted primary seawater δ O values for conodont apatite. (b) δ Oapatite trends across parts of depositional Sequences 1, 2, and 3. The relative sea-level curve is derived from sequence stratigraphic interpretations. TST = transgressive systems tract, MFZ = maximum flooding zone, HST = highstand systems tract, and SBZ = sequence boundary zone. Note that the lowest isotopic values do not coincide with the interpreted deepest water facies (MFZ) of Sequences 2 or 3, rather they occur within the shallowing 18 phase or HST, suggesting that the δ O-derived sea-level curve provides a more accurate indication of eustasy than sequence stratigraphic interpretations. See text for additional discussion. Symbols shown in Figs. 2 and 5. M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 317 318 M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 isotopic variations would preclude the occurrence of open-marine biota observed within these deposits. If the isotopic shifts were due to seawater temperature changes alone, then the amount of thermal expansion/contraction (assuming temperature-density relations of seawater to be 1.9 × 10−2 vol.%/°C) would account for b10–15 m of sea-level change, which is not sufficient to explain the observed facies juxtapositions (i.e., deep subtidal facies overlain by subaerially exposed shallow subtidal facies). The ∼ 1‰ to 3.9‰ range of up-cycle isotopic shifts for the Gray Mesa cycles are larger than that reported by Joachimski et al. (2006), but are within the range observed in the glacial–interglacial 18 Pleistocene δ O record from planktonic foraminifera (∼0.5‰ to 3.5‰; Imbrie et al., 1984). In the Pleistocene, the smaller isotopic shifts of between 0.5 and 1.0‰ shifts are related to precession-driven (∼20 ky) climate cycles, while the larger (3.5+‰) shifts reflect eccentricity-driven (∼100 ky) glacial–interglacial cycles. We suggest 18 that the wide range in δ O shifts between the two sampled Pennsylvanian cycles also reflect the difference in paleoclimate change between the various orbital frequencies. If we use the Quaternary as an analog and assume that about 30% of 18 the Pennsylvanian δ O isotopic signal is the result of tropical surface seawater temperature (SST) change and the remaining portion is due to ice-volume effects (Fairbanks and Matthews, 1978; Fairbanks, 1989), this suggests that the cycle-scale isotopic shifts of ∼1‰ to 3.9‰ represent tropical seawater temperature changes of ∼ 1.5°–6 °C, respectively, and glacio-eustatic sea-level changes of ∼55–170+ m, respectively. At the sequence-scale, the measured ∼1‰ to ∼2.4‰ variations suggest SST changes on the order of b3.5 °C and glacioeustatic changes between ∼60 and 140 m. In both cases, these are minimum estimates because deposits representing time intervals of maximum glaciation are not present at this inner to middle shelf position. Previous estimates of Middle Pennsylvanian glacio-eustatic changes derived from several different independent proxies (ice18 volume modeling, facies juxtaposition, paleotopography, and δ O isotopic shifts) range from ∼ 40 to 150 m (Adlis et al., 1988; Crowley and Baum, 1991; Heckel, 1994, 1977; Soreghan and Giles, 1999; Joachimski et al., 2006; Rygel et al., 2008). Our estimates based on 18 cycle-scale and sequence-scale δ O isotopic shifts lie within this range, and though they are minimum estimates, they suggest that the volume of Pennsylvanian glacial ice growth and melting may have been larger than the Pleistocene. Our estimated magnitudes of My sequence-scale glacio-eustatic changes are similar to or slightly smaller than those calculated for 104–105 yr changes. If this is correct, and if our interpretation that the magnitudes of sea-level change related to individual cycle development varies significantly, then this would explain why Pennsylvanian cycle-stacking patterns are so different from typical greenhouse stacking patterns. High-frequency cycles bundled within many Cambrian, Ordovician, Devonian, Jurassic, and Cretaceous greenhouse sequences are commonly characterized by thicker-than-average subtidal facies-dominated cycles developed during My-scale transgressions and thinner-thanaverage, peritidal cycles (with increased occurrence of cycle-capping exposure features) developed during My-scale regressions and lowstands (e.g., Oslger and Read, 1991; Crevello, 1991; Goldhammer et al., 1993; Elrick, 1995; Grötsch, 1996; Lehmann et al., 1998; Lehrmann and Goldhammer, 1999). Traditionally, these greenhouse stacking patterns have been attributed to relatively small magnitude, high-frequency sealevel changes superimposed on larger magnitude, My-scale sea-level changes (Goldhammer et al., 1993). In the case of Pennsylvanian icehouse cycles, the large and variable magnitude, high-frequency sea-level changes that are similar to or larger than the superimposed My-scale sea-level changes result in nonsystematic cycle-stacking patterns. For example, in the Gray Mesa Formation, some of the thinnest cycles with most intensely developed cycle-capping subaerial exposure features occur within the TST and MFZ, and some very thick cycles occur in the HST and near the accommodation minimum (Fig. 2; Scott and Elrick, 2004). Similar nonsystematic patterns of cycle thickness, intracycle facies patterns, and cycle-capping exposure features are reported for Middle and Upper Pennsylvanian cycles in the Sandia Mountains of central New Mexico (Wiberg and Smith, 1994) Orogrande Basin and Sacramento Mountains of southern New Mexico (Algeo et al., 1991; Soreghan, 1994). These characteristics suggest that cycle thicknesses and internal facies were governed by successive high-frequency accommodation trends rather than My-scale trends and contrasts with stacking patterns of typical greenhouse successions (Read et al., 1995; Lehrmann and Goldhammer, 1999; Barnett et al., 2002). 18 Of particular interest is that the sequence-scale δ O isotopic shifts do not systematically co-vary with the sequence stratigraphically derived sea-level curve (Fig. 7b). Isotopic values in the HST of sequence 2 begin to significantly decrease (implying melting of glacial ice and/or SST warming) before the facies record deepening. The 18 highest δ O values occur within the middle HST (rather than in the SBZ), and the lowest values occur within the early HST (rather than the MFZ). These trends suggest that our ability to interpret waterdepth changes from facies alone is not sensitive enough to generate 18 accurate sea-level curves. The δ O isotopic trends on the other hand, do record the environmental changes and have the potential to provide more precise indicators of glacio-eustatic changes. Additional 18 δ O isotopic data from underlying and overlying sequences at this locality and regionally coeval sections is necessary to fully understand and document the relationships between facies-derived water-depth curves and paleoclimatic and global sea-level trends. 6.4. My-scale climate drivers 18 If our δ O isotopic interpretations are correct, then it implies that a My-scale Pennsylvanian climate driver operated in tandem with highfrequency (104–105 year) climate changes to control cycle and sequence development. Over the past decade, ∼ 1–3 My-scale climatic cyclicity has been recognized in the oxygen and carbon isotope, magnetic susceptibility, foraminifera, and lithologic records of Cenozoic, Mesozoic, and Paleozoic marine successions (e.g., Lourens and Hilgen, 1997; Herbert, 1997; Shackleton et al., 1999; Matthews and Frohlich, 2002; Wade and Pälike, 2004; Coxall et al., 2005; Mitchell et al., 2008; Elrick et al., 2009). In these cases, the My-scale paleoclimate records have been attributed to long-period modulations of obliquity (∼1.2 My) and eccentricity (∼ 2.4 My) (Laskar et al., 1993; Laskar, 1999; Laskar et al., 2004). We suggest that the My-scale paleoclimate changes in the Middle Pennsylvanian were driven by long-period obliquity variations which lead to glacial ice growth and melting in southern Gondwana, My-scale glacio-eustatic sea-level oscillations with magnitudes of many tens to over a hundred meters, and the development of globally widespread 3rd-order depositional sequences. 7. Conclusions 1) Deep subtidal through shallow subtidal facies of the Middle Pennsylvanian Gray Mesa Formation are grouped into ∼75 highfrequency (104–105 yr) upward-shallowing subtidal cycles. Over 50% of the cycles are capped by early diagenetic features including pedogenic calcretes, regolith intraclasts, and diagenetic mottling. The cycles are stacked into 4 1/2 My-scale (3rd-order) depositional sequences (40–80 m). 13 2) High-resolution δ C stratigraphy across nine of the cycles indicates that the cycle tops were diagenetically altered by isotopically light meteoric fluids during high-frequency sea-level fall and lowstand. The base of the cycles retain marine or near 13 marine δ C values due to diagenesis in rock-dominated pore M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320 13 fluids. Cycle-capping early diagenetic features preserve δ C isotopic evidence of substantial meteoric diagenesis during 13 subaerial exposure. These combined δ C isotopic patterns demonstrate that sea-level changes related to cycle development can be detected even within carbonate successions displaying subtle facies changes and/or development of subaerial exposure features. 18 3) δ O values of conodont apatite from two sampled cycles support the hypothesis that the high-frequency sea-level changes were driven 18 by glacio-eustasy combined with changes in SST. δ O values from conodont apatite spanning parts of three depositional sequences also indicates that My-scale (3rd-order) glacio-eustasy and SST changes controlled sequence development. The difference between lithologically interpreted sea-level curves and those determined 18 from δ O trends suggests that oxygen isotopes are a more sensitive indicator of glacio-eustatic changes. Using the Pleistocene icehouse as a modern analog, estimates on the magnitudes of glacio-eustasy for the sampled cycles are between ∼55 and 170+ m combined with tropical SST changes of ∼1.5°–6 °C. Estimated sequence-scale glacio-eustatic oscillations range from ∼60 to 140 m with SST changes of b3.5 °C. The most plausible driver for the My-scale paleoclimate changes is long-period obliquity (∼1.2 My) variations. 4) The calculated high-frequency glacio-eustatic values are similar or greater than Pleistocene values, and lie within the range estimated for other Middle Pennsylvanian successions using a variety of independent eustatic proxies. 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