Impacts of basin restriction on geochemistry and extinction patterns: A case from the Guadalupian Delaware Basin, USA

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.1 (1-11)

Earth and Planetary Science Letters ••• (••••) ••••••

Contents lists available at ScienceDirect

Earth and Planetary Science Letters www.elsevier.com/locate/epsl

Impacts of basin restriction on geochemistry and extinction patterns: A case from the Guadalupian Delaware Basin, USA Benjamin P. Smith a,∗ , Toti Larson b , Rowan C. Martindale a , Charles Kerans a,b a

Jackson School of Geosciences, Department of Geological Sciences, The University of Texas at Austin, Austin, TX 78712, USA Bureau of Economic Geology, John A. and Catherine G. Jackson School of Geosciences, University of Texas at Austin, 10100 Bureau Road, Building 130, Austin, TX 78713-8924, USA b

a r t i c l e

i n f o

Article history: Received 12 December 2018 Received in revised form 12 August 2019 Accepted 3 October 2019 Available online xxxx Editor: L. Derry Keywords: carbonate geochemistry carbon isotope chemostratigraphy restricted basin Guadalupian Mid-Capitanian extinction Delaware Basin

a b s t r a c t Geochemical data from carbonates often constrain the nature of environmental change during biotic turnover events. Many ancient carbonates, however, formed in geographically-isolated basins subject to local environmental factors, resulting in varying extinction rates between open ocean and restricted settings. It follows that high-resolution data from restricted basins may help unravel poorly-understood biotic crises such as the Mid-Capitanian extinction, which had especially high extinction rates in restricted settings. This study examines factors controlling salinity, stratification, and oxygenation in the Capitanian (Middle Permian) Delaware Basin, USA. Elemental and carbon isotope measurements from time-equivalent strata reveal differences between shallow- and deep-water masses, pointing to local controls such as stratification and de-oxygenated bottom water. Basinal dolomites and evaporites mark periods of elevated salinity tied to sea-level lowstands, which correspond with turnovers in fusulinid and brachiopod communities. Faunal turnover in the Delaware Basin demonstrates a fundamental attribute of restricted basins: water chemistry is often tightly coupled to physical process such as sea level change. We suggest that the relationships among sea level fluctuations, chemical changes, and biotic turnover may explain why the Capitanian mass extinction was more severe in isolated basins than the open ocean. © 2019 Elsevier B.V. All rights reserved.

1. Introduction Carbon isotopes and trace elements in carbonates can record interactions between marine biota and their environments (Kump and Arthur, 1999; Payne and Kump, 2007). Many biotic crises are associated with physical and chemical restructuring of the ocean, including changes in the oxygen minimum zone (Lau et al., 2016), thermohaline circulation (Saltzman, 2003), and the availability of surface nutrients (Meyer et al., 2011). Such changes may manifest in the chemical composition of carbonates because they are sensitive to vertical gradients in temperature and redox conditions (Immenhauser et al., 2002; Hood and Wallace, 2015). Examined together, data from coeval shallow- and deepwater carbonates can reveal first-order changes in the structure of the water column, constraining the relationship between biotic turnover and paleoceanographic change (Meyer et al., 2011; Song et al., 2013).

*

Corresponding author at: Jackson School of Geosciences, 2305 Speedway Stop C9000, Austin, TX 78712-1692, USA. E-mail address: [email protected] (B.P. Smith). https://doi.org/10.1016/j.epsl.2019.115876 0012-821X/© 2019 Elsevier B.V. All rights reserved.

Carbonate geochemistry, however, does not uniquely preserve global environmental conditions. In addition to diagenetic resetting, local factors may obscure or even supersede global-scale patterns. Local factors are especially important because many ancient carbonate platforms developed in geographically isolated basins or interior seaways (Fig. 1A). Geochemical trends from isolated basins often depart from global trends because they are more sensitive to changes in freshwater input, terrigenous input, circulation, and bottom water oxygenation (Holmden et al., 1998; Immenhauser et al., 2008). Another local factor to consider is downslope transport, which re-distributes sediments along the platform profile. If shallow- and deep-water sediments fix different geochemical properties from seawater—or experience different conditions during marine diagenesis—then shifts in downslope transport patterns may be mistaken for palaeoceanographic change (Swart and Eberli, 2005). Thus, any use of carbonate geochemistry for paleoceanographic interpretation must be considered within the regional stratigraphic context. Despite local complications, isolated basins should be studied, rather than dismissed. Basin connectivity or restriction can drive larger patterns of origination and extinction (Stigall, 2010). Biotic Immigration Events during sea level highs, in tandem with vi-

JID:EPSL

2

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.2 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

Fig. 1. Geologic setting of the study area after Scotese and Langford (1995) and Warren (2010). A) Permian epieric seaways and evaporite deposits along the margins of Pangea: 1. Delaware Basin (USA), 2. Phosphoria seaway (Western USA), 3. North American Midcontinent (USA) 4. Zechstein Basin (Europe), 5 Pre-Caspian depression and Uralian Foredeep (Russia and Khazakstan), 6. East Greenland seaway, 7. Khuff Formation (Arabian Plate). B) Paleogeographic map of the Delaware Basin during the Middle Permian after Tinker (1998). By the Capitanian (Middle Permian), the basin was restricted; the Hovey Channel was the only source of exchange with the Panthalassa Ocean. At the end of the Middle Permian, carbonate platforms associated with the Capitan Reef died off and the basin was filled with deep-water evaporites.

cariance (the geographical separation of basins) during sea level lows, are hypothesized to have promoted speciation during the Great Ordovician Biodiversification Event (Stigall et al., 2017). In contrast, during the Frasnian-transgressive pulses, interbasinal biotic invasions are thought to have decreased speciation events by shutting off vicariance (speciation due to the physical separation of a group by a geographical barrier) and ultimately contributing to depressed diversity during the late Devonian biotic crisis (Stigall, 2010). Furthermore, isolated basins may have distinct records of biotic events. For example, per-capita extinction rates for the mid-Capitanian crisis are significantly higher in isolated settings than the open ocean, suggesting an amplification of the extinction trigger(s) in epicratonic basins (Miller and Foote, 2009). Building from this idea, we propose that isolated marine basins chemically modify global environmental signals because physical processes, especially sea-level change or basin closure, trigger corresponding chemical changes such as anoxia or elevated salinity. Documenting local water mass evolution with carbonate geochemistry serves as a stepping stone to unravelling larger-scale relationships between biotic turnover and environmental changes. This study investigates links among relative sea level change, local water mass evolution, and the paleontological record in the Delaware Basin, USA, where basin closure coincided with the Capitanian extinction and biodiversity crisis. The Capitan Reef and its associated platforms provide numerous shelf-to-basin profiles spanning the entire Capitanian Stage (Kerans and Kempter, 2002). We employ several approaches to constrain the physical and chemical evolution of the Delaware Basin. First, we use petrographic evidence for early dolomite and evaporite minerals to constrain periods of elevated salinity. Second, we compare δ 13 C values from the Delaware Basin and the open ocean to identify the presence of stratified or poorly-circulated water (e.g., Holmden et al., 1998; Song et al., 2013). Third, we use redox-sensitive trace elements (manganese, vanadium, and uranium) in basinal sediments as proxies for bottom-water de-oxygenation (e.g., Algeo and Maynard, 2004; Schroder and Grotzinger, 2007). We compare geochemical data with previous paleontological studies to determine which changes, if any, have the strongest effect on local biota. Finally, we extend the implications of our results towards understanding larger-scale patterns in Capitanian extinction rates.

2. Geologic background 2.1. The Mid-Capitanian extinction and ecological crisis The Capitanian Stage hosted an important mass extinction and biodiversity crisis, though there is no consensus on the severity and cause(s) (Bambach et al., 2004; Clapham and Payne, 2011). Proposed extinction mechanisms include sea level fall (Hallam and Wignall, 1999), global cooling (Isozaki et al., 2007), marine anoxia (Clapham et al., 2009), and toxic metal poisoning (Grasby et al., 2016). Many potential kill mechanisms have been linked to Emeishan volcanism in China based on a −6h δ 13 C excursion (Bond et al., 2010; Zhou et al., 2002). Nevertheless, recent carbon and calcium isotope work shows no globally correlative Mid-Capitanian excursion, suggesting only a small effect on the global carbon cycle (Jost et al., 2014). One clear pattern, however, is that the Capitanian extinction was more severe in epicratonic settings than the open ocean (Miller and Foote, 2009). Detailed studies of restricted settings like the Delaware Basin may uncover relationships that are useful for interpreting larger-scale extinction patterns. 2.2. Palaeoceanographic records in carbonates Geochemical methods for paleoceanography take advantage of relationships among organic carbon cycling, redox conditions, and vertical mixing. In modern oceans, the balance between vertical mixing and biological pumping creates a relatively small (<3h ) δ 13 C gradient (Song et al., 2013). Larger δ 13 C gradients require strong density stratification or high surface productivity. In the Black Sea, strong density stratification sustains an ∼8h δ 13 C gradient between surface and deep waters (Fry et al., 1991). Alternatively, high surface productivity has been invoked to explain >3h δ 13 C gradients in the geologic record (Meyer et al., 2011). In both cases, re-mineralization of organic matter will likely deplete available oxygen in the water column. Redox-sensitive trace elements—e.g., manganese, vanadium, and uranium—can be enriched or depleted in sediments according to the availability of oxygen, and suites of these elements can be used to infer changes in bottom-water oxygenation (Algeo and Maynard, 2004; Tribovillard et al., 2006). When preserved, redox-indicators and large δ 13 C gradients place important constraints on regional and global paleoceanography.

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.3 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

3

Fig. 2. Stratigraphic nomenclature for shelfal, reefal, and basinal units in the Delaware Basin. Sequence picks and correlations follow Kerans and Kempter (2002) and conodont zones from Lambert et al. (2002) are shown in grey text. Word. = Wordian, Wuch. = Wuchiapingian, J. = Jinogondolella.

2.3. Stratigraphic and palaeoceanographic setting of Delaware Basin samples The Delaware Basin formed during the late Paleozoic, when continental collisions created several restricted basins and interior seaways across North America (Yang and Dorobek, 1995). By the early Capitanian Stage (Middle Permian), the Delaware Basin was rimmed by steep-walled platforms associated with the Capitan Reef complex (Fig. 1). The basin was likely connected to the Panthalassa Ocean through one or more southern inlets such as the Hovey Channel (Fig. 1B). At the end of the Capitanian Stage, carbonate platforms completely infilled the channels, and the isolated basin filled with evaporites (Garber et al., 1989). The sedimentary history, which culminates in basin-wide evaporite deposits, suggests that Delaware Basin water may have evolved substantially during the Capitanian Stage. Shelf-to-basin outcrops in the Guadalupe and Apache Mountains present an excellent opportunity to test for time-dependent paleo-depth gradients using carbonate geochemistry. We take advantage of previous work in sequence stratigraphy, biostratigraphy, and depositional systems to compare time-equivalent shelfal and basinal deposits (Kerans and Kempter, 2002; Nestell and Nestell, 2006). Because the mid-Capitanian extinction was especially severe for fusulineacean foraminifera and brachiopods (Bond et al., 2010; Shen and Shi, 1996), we also integrate work on these taxa in the Delaware Basin (Fall and Olszewski, 2010; Olszewski and Erwin, 2009; Wilde et al., 1999). These studies provide intra-Capitanian resolution through use of a regional sequence stratigraphic framework and conodont biozones (Lambert et al., 2002). Conodont biozones also allow comparison among basins to determine whether geochemical changes are global or local. 3. Methods 3.1. Sample collection and preparation Samples from shallow- and deep-water environments were collected from the Guadalupe and Apache Mountains (supplemental Table S1). Shallow-water samples come from paleo-aragonite cements in supratidal environments; we prioritized these samples because previous geochemical work concluded that cements underwent diagenesis in low water-to-rock conditions and likely have primary δ 13 C values (Chafetz et al., 2008). Deep-water samples come from carbonate-bearing members of the basinal Cherry Canyon and Bell Canyon Formations, which were deposited in the

toe-of-slope and basin environments with a maximum water depth of 400–600 m (Tinker, 1998). Basinal carbonates are vertically separated by thick silicilastic packages, which preclude collection of continuous geochemical profiles. Instead, we grouped available stratigraphic, paleontological, and geochemical data into five Wordian to Capitanian subdivisions (Fig. 2). From oldest to youngest, these groupings are 1) the basinal Manzanita Member, which sits below the basal Capitanian Stage boundary; 2) the Seven Rivers platform, which includes the Seven Rivers Formation and the basinal Hegler and Pinery Members; 3) the Yates Platform, which includes the terminal Yates shelf top strata and the McCombs and deep-water McKittrick limestones; 4) the lower Tansill platforms, which include Tansill strata below the Ocotillo siltstone and the basinal Lamar limestone; and 5) the upper Tansill, which includes the Tansill Formation above the Ocotillo siltstone and the basinal Reef Trail member. Hand samples were cut into thin section billets and stained with Alizarin Red S to discriminate between calcite and dolomite. Thin sections were described and visually screened to select samples with low porosity and primarily inclusion-rich (likely marine) cements. Whenever possible, we avoided fractures, stylolites, clear cements, and recrystallized skeletal material. Billets were sampled with a one-millimeter tungsten carbide-tipped dental drill, and stable isotope and trace element analyses were performed on splits of a single homogenized powder to minimize the effects of intersample heterogeneity. 3.2. Carbon and oxygen isotopes All δ 18 O and δ 13 C analyses were performed on a Finnigan MAT253 coupled to a ThermoFischer Gas Bench II at the University of Texas at Austin. For each sample, splits of 250–700 μg powders were weighed and dried overnight at 70 ◦ C. Vials were sealed and flushed with helium gas before samples were reacted with phosphoric acid (H3 PO4 ) for two hours at 50 ◦ C. Analytical precision (1σ ) for all runs was better than ±0.09h for δ 13 C and ±0.19h for δ 18 O based on replicate analyses of internal lab standards. Measurements were calibrated to NBS18 and NBS19 international standards and reported relative to V-PDB. Data analysis for carbon isotopes consisted of two kinds of comparisons. First, we grouped δ 13 C values by stratigraphic interval and compared each group to a reference value for seawater. We selected brachiopod data from Korte et al. (2005) as a reference for global seawater δ 13 C because the dataset has good spatial and temporal coverage for the Capitanian Stage. We also compared

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.4 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

4

Fig. 3. Cross-plots of δ 18 O, δ 13 C, and Mn/Sr data. Screened-data are those removed using the δ 18 O filter described in Section 4.1. Solid lines are analytic solutions for fluid-rock interaction in an open system as given by Jacobsen and Kaufman (1999) with parameters from their figures 1 and 2 except where noted. The initial limestone composition is modeled as Mn = 1 ppm, Sr = 100 ppm with estimated original δ 18 O = +2h and δ 13 C = +8h based on the highest values in the data; initial fluid composition is set at Mn = 4 ppm, Sr = 50 ppm, C = 1000 ppm with combined fluid isotope composition and fractionation factors of −14h for oxygen (δ 18 Ow0 + O ) and −12h for carbon (δ 13 Cw0 + C ).

time-equivalent shallow-water and deep-water δ 13 C values to see if there were statistically-significant differences. All comparisons were carried out using one- and two-sample Wilcoxon signed-rank tests (wilcox.test function) in R (version 3.4.3, R Core Team, 2017) at the 95% confidence level with a Bonferroni correction to account for multiple comparisons (n = 12). 3.3. Elemental concentrations Elemental concentrations were measured at the University of Texas at Austin with an Agilent 7500ce quadrupole ICP-MS running in solution mode. For each sample, 10 mg splits of powder were acidified with 0.33 N acetic acid (CH3 COOH), then agitated by inclined rotation overnight. The acid volume was 80% of the molar volume required to dissolve a sample of pure carbonate, minimizing reactions with any non-carbonate phases (Etemad-Saeed et al., 2016). We note that while our method minimized contamination from clay minerals, it may exclude redox indicators that associated with non-carbonate phases; [Mo] and [Fe] were measured but were generally below detection in most samples. Leachates were centrifuged and diluted with 2% nitric acid (HNO3 ) for analysis. Uncertainties for Mg, Al, Ca, V, Mn, Sr, and U were <7% based on replicate analyses of both in-house standards and the NIST1643f international standard. The relative standard deviation for replicate sample analyses was used as a measure of in-run precision; the average relative standard deviation for Mg, Ca, Mn, and Sr was <5% across all samples while the average relative standard deviation for Al, V, and U was <7% for basinal samples. Clay and detrital minerals contain high concentrations of redoxsensitive trace elements, and even minor amounts of these phases can mask palaeoceanographic trends (Algeo and Maynard, 2004). Following the approach of Tribovillard et al. (2006), the authigenic fraction of each redox indicator (Xauth ) was estimated as:



Xauth = Xsample −

Xshale Alshale



∗ Alsample

where Xshale and Alshale are the concentrations of each redox indicator and aluminum in post-Archean Average Shale. The authigenic concentrations were then normalized to Ca, and significant trends in X/Ca ratios were determined by one-way analysis of variance (ANOVA, aov function in R) (R Core Team, 2017). Pairwise comparison of means between intervals was conducted using Dunnett’s modified Tukey-Kramer post hoc test, which can handle unequal variances (e.g., Lau et al., 2016).

4. Results 4.1. Geochemical indicators of diagenetic overprint We used δ 18 O, δ 13 C, and Mn/Sr ratios to screen samples for evidence of late fluid-rock interaction (Fig. 3). Data were first screened according to oxygen isotopes because marine carbonates typically acquire low δ 18 O values during meteoric or burial diagenesis. For each formation, only samples above a cutoff value—one standard deviation below the mean—were considered for further analysis of carbon isotopes and redox indicators (sensu Metzger and Fike, 2013). Covariation among δ 18 O, δ 13 C, and Mn/Sr was then calculated from the screened data, excluding dolomitic samples. We also employed a simple diagenesis model (Jacobsen and Kaufman, 1999) to show generalized covariation patterns expected under progressive fluid-rock interaction with either burial fluids and meteoric water. Correlations between δ 18 O–δ 13 C and Mn/Sr – δ 18 O were significant, but only weakly positive (cor.test function in R, Spearman’s ρ < 0.2, p < 0.05) while Mn/Sr-δ 13 C displayed a moderately negative correlation (Spearman’s ρ < −0.4, p < 0.001). 4.2. Facies and dolomitization in basinal carbonates Observed textures in basinal carbonates are consistent with deposition by sediment gravity flows as documented by Brown and Loucks (1993). Beds comprise decimeter to meter-scale packages with Dunham textures ranging from laminated wackestones to intraclast-skeletal grainstones and rudstones. Common sedimentary structures include normal grading and rip-up clasts. Fauna consists of a mix of shelfal fossils (e.g., fusulinids) as well as Tubiphytes and brachiopods sourced from the reef and slope. The basinal members all shared a similar assortment of facies, although the relative facies proportions varied. Diagenetic textures varied widely among basinal deposits with certain intervals almost entirely replaced by dolomite and evaporite minerals. Just below the Wordian/Capitanian boundary, the lower half of the Manzanita carbonate consists of fine-grained dolomite. The Manzanita Mb. also contains centimeter-sized bedding-parallel vugs interpreted by Hampton (1989) as evaporite molds (Fig. 4E). In the upper Manzanita Mb., lath-shaped anhydrite pseudomorphs are common in thin sections above the dolomitelimestone contact. These molds preferentially replace Tubiphytes and predate inclusion-free calcite spar (Fig. 4A). In the midCapitanian, the Yates-equivalent basinal deposits contain fabric-

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.5 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

5

Fig. 4. Basinal dolomite and evaporite textures in the Delaware Basin. A) Lath-shaped calcite pseudomorphs after anhydrite replacing Tubiphytes in the Manzanita Member (latest Wordian). B) Evaporite molds in the Lamar Member (Jinogondolella postserrata/altudensis transition). Notice the compaction of mm-scale bedding around the mold, indicating an early timing for evaporite growth. C) Partial dolomite replacement in the McCombs Member (Jinogondolella postserrata zone). Dolomite replacement preferrentially occurs in the fine-grained tops of dilute turbidites (pictured) and in the matrix of poorly-sorted debris flows. D) Calcitized pseudomorphs after gypsum (?) in the Reef Trail Member (Jinogondolella altudensis zone). E) Vuggy weathering (interpreted as evaporite molds) in the dolomitic portion of the Manzanita Member. Tu = Tubiphytes, Ev = evaporite mold.

destructive, fine-grained dolomite. Dolomite replacement is fabric selective, with individual fining-upwards beds often showing a change from pure limestone to >90% dolomite over a few centimeters (Fig. 4C). In poorly-sorted beds, dolomite is concentrated in the matrix while larger skeletal grains are still calcitic. Both textures suggest that dolomite preferentially replaced carbonate mud. No evaporite textures were found in association with the Yatesequivalent basinal limestones. Thin section evidence suggests that evaporite growth occurred early in the diagenetic history rather than during late burial diagenesis. In the Tansill-equivalent basinal deposits, lath-shaped anhydrite pseudomorphs are preserved along bedding planes (Fig. 4B). Thin laminae in mudstones and wackestones show pressure halos around evaporite molds, demonstrating that evaporite growth predated significant compaction. Evaporites must have grown early in the burial history because carbonate mud loses ∼50% of its volume under 150 m of overburden (Goldhammer, 1997). Basinal evaporite textures are distinct from late generations of pore-filling anhydrite that are common in the Delaware Basin subsurface (Garber et al., 1989). 4.3. Redox-sensitive trace elements in basinal carbonates Redox indicators (Mn/Ca, V/Ca, and U/Ca) vary by several orders of magnitude when grouped by stratigraphic age (Fig. 5). To guide interpretations, we modeled expected values as:

(X/Ca)carbonate = K d ∗ (X/Ca)seawater

where X/Ca denotes the metal/Ca ratio and K d is the effective distribution coefficient. While simple, this model applies to both equilibrium and non-equilibrium conditions (Jacobsen and Kaufman, 1999). The model uses the fluid composition of modern seawater and a range of experimentally-determined distribution coefficients in seawater-like solutions (supplemental Table S2). The spread in modeled X/Ca ratios represent kinetic variations due to differences in minerology, precipitation rate, and vital effects among the surveyed studies. Enrichments are defined as values above the highest modeled values. Manganese concentrations show mean enrichments of ∼2–4 orders of magnitude above modeled values for all time intervals. The highest values are found in oldest deposits (mean Mn/Ca = 2.3 × 10−3 ) and steadily decrease to a mean of 2.4 × 10−4 for the youngest deposits. Vanadium concentrations are only slightly enriched in two oldest platforms (Seven Rivers and Yates equivalents, means from 4.4 × 10−7 to 1.0 × 10−6 ) but show higher enrichments up to 2 orders of magnitude for the two youngest basinal deposits (lower and upper Tansill equivalents, means from 6.0 × 10−6 to 1.5 × 10−5 ). U/Ca means range from 1.2 × 10−5 to 4.2 × 10−5 and fall within the range of modeled values. Oneway ANOVA models applied to trace element data indicate that the means for Mn/Ca, V/Ca, and U/Ca all have significant variations at the 95% confidence interval (supplemental Tables S3–S5). Post hoc tests the lower and upper Tansill platforms have lower mean Mn/Ca ratios and higher mean V/Ca ratios compared to initial val-

JID:EPSL

6

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.6 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

Fig. 5. Trace metal/calcium ratios in basinal carbonate members grouped by stratigraphic age. Yellow boxes indicate modeled ranges using the composition of modern seawater and experimentally-determined distribution coefficients (supplementary Table S2). One-way analysis of variance (ANOVA) models for all three elements indicate that the means are not identical across all four time intervals at the 95% confidence level. An asterisk (∗) marks intervals that have significantly different means from the initial mean (Seven Rivers). Significance was determined by a Dunnett’s modified Tukey-Kramer post-hoc test (supplementary Tables S3–S5). Mn/Ca values are elevated across all intervals but decreases steadily throughout the Capitanian. In contrast, V/Ca is enriched across the Yates-lower Tansill boundary. U/Ca shows small oscillations throughout the Seven Rivers to lower Tansill but shows no significant enrichments above modeled values. (For interpretation of the colors in the figure(s), the reader is referred to the web version of this article.)

Fig. 6. Comparison of bulk rock carbon isotope values in shallow-water (red) and deep-water (blue) depositional environments throughout the Capitanian. The gray lines represent globally-averaged δ 13 C values from unaltered Permian brachiopods (Korte et al., 2005). A single asterisk (∗) denotes values that are distinct from time-equivalent seawater as determined by a one sample Wilcoxon signed-rank test (supplementary Table S6). Two asterisks (∗∗) denote time-equivalent shallow- and deep-water samples whose means are distinct from one another as determined by a two-sample Wilcoxon signed-rank.

ues at the beginning of the Capitanian Stage (Seven Rivers). U/Ca values, however, show no overall trends. 4.4. Spatial and temporal patterns in carbon isotopes When grouped by stratigraphic age, mean δ 13 C values are consistently greater for shallow-water strata than time-equivalent deep-water environments (Fig. 6). For shallow-water environments, mean carbon isotope values increase steadily from +4.4h to +6.5h throughout the Capitanian Stage. Mean deep-water values are more variable, showing an increase from +3.0h to +5.4h across the Seven-Rivers to Yates platforms, followed by a decrease to +2.7h for the late Tansill. A two-sample Wilcoxon signed-rank test assessed the significance of any offset between shallow-todeep pairs (δ 13 C). All time intervals except for the Yates dis-

played significant differences in means, with a maximum offset of 3.9h in the upper Tansill. To constrain overall deviation from global seawater values seawater, a one-sample Wilcoxon signed-rank test was used to compare the mean of each group to a globally-averaged values reported by Korte et al. (2005). Six of the eight stratigraphic units have mean δ 13 C values that are significantly different from global reference values (Fig. 6, supplemental Table S6). All shallow-water units except the Seven Rivers show positive deviations from seawater, while all deep-water units except the Yates-equivalent and lower Tansill-equivalent show significant negative deviations. For the lower and upper Tansill, neither shallow- nor deep-water samples give an adequate approximation of seawater δ 13 C values. Given the apparent disagreement of these δ 13 C data with values from Korte et al. (2005), we consider whether choosing different reference values for Permian seawater would change the result.

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.7 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

Higher δ 13 C values have been reported for the latest Capitanian Stage; for example, Isozaki et al. (2007) reported values as high as +6.0h from southwest Japan. Even assuming a seawater value as high as +6.0h does not account for +6.5h average for the upper Tansill shallow-water samples (p < 0.001). The maximum δ 13 C measurements, which are typically interpreted as the least altered, reach values as high as 8.6h for the upper Tansill Fm. Lastly, a higher reference value for seawater fails to address the 3.9h offset between shallow- and deep-water samples in the upper Tansill Fm. Overall, the Capitanian δ 13 C values reported here are difficult to wholly reconcile with previously published δ 13 C values believed to be representative of global seawater. 5. Discussion The δ 18 O, δ 13 C, and Mn/Sr data place important constraints on the diagenetic history and, by extension, the suitability of samples for paleoceanographic interpretation. The most important question is whether the geochemical patterns have been entirely reset by either late burial or uplift-related meteoric diagenesis. The screened δ 18 O–δ 13 C data (Fig. 3A) do not display a strongly hyperbolic curve predicted by most fluid-rock interaction models, although more linear arrays can develop through interaction with high pCO2 burial fluids (Derry, 2010). Likewise, the Mn/Sr–δ 13 C trend alone does not preclude either a burial or meteoric diagenetic history (Fig. 3B). The Mn/Sr–δ 18 O patterns (Fig. 3C), however, are inconsistent with late meteoric or burial diagenesis; the screened data show a positive correlation with the highest Mn/Sr ratios (>40) associated with high δ 18 O values (>0h). In contrast, high Mn/Sr ratios that develop during meteoric or burial diagenesis would be associated with negative δ 18 O values as illustrated by the Jacobsen and Kaufman (1999) model. While it is unlikely that the screening procedure completely removed later diagenetic overprint, the screened Mn/Sr–δ 18 O data suggests primary seawater values modified by fluid-rock interaction during early marine or marine burial diagenesis (sensu Melim et al., 2001). Given the geologic history, it is reasonable to ask whether basin closure might cause secular changes in both primary seawater composition and the marine burial environment. Below, we summarize the petrographic and geochemical evidence for changes in salinity, stratification, and bottom-water oxygenation. The timing of these changes is critical for interpreting the local paleontological record and lends insight into environmental stresses in restricted marine settings overall. 5.1. Basinal dolomite and evaporite molds as markers of basinwide salinity change Early dolomitization in the Delaware Basin is among the most obvious and volumetrically significant evidence of diagenesis with seawater-derived fluids. The Capitan platforms had large, evaporitic lagoons that drove reflux dolomitization, resulting in dolomitic shallow-water deposits that grade into limestones in the basin (Melim and Scholle, 2002). The only notable exceptions are the dolomites in the Manzanita and the Yates-equivalent basinal members (Section 4.2). Furthermore, basinal dolomites are texturally distinct from shallow-water dolomites because they preferentially replace lime mud (Fig. 4C) rather than high-permeability fabrics such as grainstones and fractures in the reef (Melim and Scholle, 1999). The spatial distribution and diagenetic textures suggest that while basinal dolomites formed early, they are nevertheless distinct from the dominant generation of reflux dolomite. Stratigraphic correlation of basinal dolomites suggests an alternative model for their origin. The Manzanita Member correlates with a prominent exposure surface on the shelf marked by siliciclastic bypass (Kerans and Kempter, 2002). Likewise, the

7

Yates-equivalent basinal members bracket another unconformity associated with paleokarst and bypass (Kosa and Hunt, 2006). We propose a model in which extended sea level lowstands decreased inflow through the channel(s) connecting the basin to the ocean, elevating salinity to the point where dolomite and rare evaporite minerals precipitated (Fig. 6B). This model is supported by petrographic evidence that evaporite textures formed before significant burial (Section 4.2). Although only minor volumes of dolomite and evaporites precipitated during lowstands, they provide an important benchmark for the chemical evolution of basin water through time. The Capitanian interval is bracketed below by basinal dolomite and evaporite molds (Manzanita member) and above by >500 m of evaporites (Castile Formation). The isolated water mass of the Delaware Basin experienced periods of elevated salinity throughout the entire Capitanian and not just near the end of the Capitanian Stage. 5.2. Redox-sensitive trace elements as indicators of bottom-water de-oxygenation Both V/Ca and Mn/Ca data display statistically-significant shifts through time. Additionally, the highest recorded values are several orders of magnitude above modeled baselines. For modeled values, either the assumed fluid ratios are too low, or else the effective distribution coefficients are too small (Fig. 5). High values might result from higher proportions of Mn- and V-bearing compounds, which in turn affects the availability of these elements during carbonate diagenesis (Algeo and Maynard, 2004; Tribovillard et al., 2006). Thus, the Mn/Ca and V/Ca shifts likely represent a combination of changes in the delivery of these elements to sediments as well as changes in redox conditions during early diagenesis. The opposing V/Ca and Mn/Ca shifts can be explained by decreasing bottom-water oxygenation throughout the Capitanian Stage. Early Capitan platforms have comparatively low V/Ca ratios and high Mn/Ca ratios. In the modern ocean, Mn-rich carbonates form when the overlying water column in oxic, which promotes Mn export to sediments as solid (oxy)-hydroxide phases (Algeo and Maynard, 2004; Tribovillard et al., 2006). In contrast, low-oxygen sediments display slightly lower Mn enrichments as some Mn is lost through diffusion (Calvert and Pederson, 1993). Vanadium enrichment begins when bottom-water becomes dysoxic (Algeo and Maynard, 2004), and the increase in V/Ca between the Yates and Tansill platforms represents the first appearance of such conditions. These data agree with paleontological work which also suggests at least intermittent dysoxic conditions on the lower Tansill slope (Babcock, 1977). Interestingly, uranium enrichment should accompany vanadium enrichment, but no significant U/Ca trends were observed. Either U enrichments were reset during post-depositional re-oxygenation (Lau et al., 2017; Tribovillard et al., 2006) or uranium preferentially entered a noncarbonate phase, which would not be detected in our method (section 3.3). The combined redox and salinity histories provide a detailed picture of chemical changes in the Delaware Basin (Fig. 7). Decreased exchange with the open ocean not only results in periods of elevated salinity but also bottom water de-oxygenation. Similar dynamics have been observed in basins of the southern California borderland (e.g., Berelson, 1991); although still connected to the Pacific Ocean, basins are barred by topographic highs (sills) and frequently experience restriction and stagnation, including deoxygenation of bottom waters. Furthermore, experimental work with brine chemistry suggests an underlying relationship between elevated salinity and dissolved oxygen: at high ionic strengths, dissolved gasses are “salted out” of solution. At 25 ◦ C, there is a 25% reduction in dissolved oxygen as salinity increases from 35h to

JID:EPSL

8

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.8 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

Fig. 7. Evolution of the Delaware Basin water mass throughout the Capitanian Stage. A) In the earliest Capitanian, the Delaware Basin was mostly well oxygenated and open marine. B) A prolonged sea level fall event near the mid-Capitanian resulted in an increase in basin salinity and partial replacement of basinal lime muds by dolomite. Redox-sensitive trace elements indicate the basin was still well-oxygenated. C) Late Capitanian platforms are characterized by stagnant, deoxygenated bottom water and large vertical δ 13 C gradients due to stable salinity stratification.

the onset of gypsum precipitation at ∼72h (Debelius et al., 2009). This may explain why bottom-water de-oxygenation in the Tansill platforms directly follows an increase in salinity in the Yates platform (Fig. 7). Taken together, the co-evolution of salinity and oxygen in the Delaware Basin demonstrates how physical changes in oceanic exchange can trigger a cascade of related chemical changes in restricted basins. 5.3. Combined sedimentological and environmental influences on shelf-to-basin δ 13 C patterns The anomalous δ 13 C record in section 4.4 requires an explanation for departures from accepted Permian seawater values. The low δ 13 C values of basinal carbonates relative to time-equivalent seawater could represent diagenetic exchange with isotopicallynegative, organic-derived carbon. This process was likely limited, though, because basinal carbonates typically have <0.4% total organic carbon (Jin et al., 2012). More importantly, diagenetic alteration with organic-derived carbon cannot explain why shallowwater values are consistently heavier than time-equivalent seawater, up to a maximum of 8.6h in the upper Tansill platform. At least one additional driver must be invoked. High δ 13 C values could be due to either variations in original minerology or unusual oceanographic conditions. Under similar conditions, the δ 13 C of aragonite is up to ∼1.7h heavier than calcite (Romanek et al., 1992). Another possibility is that biological pumping leaves surface waters with residually higher δ 13 C values, up to 3h in modern open ocean settings (Song et al., 2013 and references therein). Since the shallow-water samples have an

aragonitic precursor, either of these processes could be responsible for the 1.4h offset between shallow- and deep-water samples in the Seven Rivers and lower Tansill platforms. In contrast, the 3.9h δ 13 C gradient in the lower Tansill platform is too large to explain with either of these processes. A >3h δ 13 C gradients can result from enhanced primary productivity (Meyer et al., 2011) or severe water column stratification (Song et al., 2013). We favor stratification as the likely mechanism because the largest gradient precedes deposition of deep-water evaporites, and because the upper Tansill has comparatively less siliciclastic material than underlying platforms, suggesting less nutrient influx into the basin. The above discussion illustrates some of the challenges in interpreting shelf-to-basin δ 13 C patterns in steep-walled platforms. If global carbon cycle changes control δ 13 C variability, then shallowand deep-water records should mimic one another with only a minimal (<3h) offset (Fig. 8A). The changing offset between shallow- and deep-water δ 13 C in Fig. 5 indicates that there is at least one additional control. Two additional processes can generate δ 13 C offsets: changes in downslope transport (e.g., Swart and Eberli, 2005) and changes in environmental gradients (e.g., Song et al., 2013). Changes in downslope transport are extremely difficult to constrain, but such processes only serve to reduce the apparent δ 13 C gradient by homogenizing sediment; any observed gradient should be considered a minimum estimate for the actual seawater gradient (Fig. 8B). The development of large (>3h δ 13 C) offsets between shallow- and deep-water cannot be attributed to sediment transport alone. Thus, the appearance of large δ 13 C gradients can be compelling evidence for paleoceanographic changes, especially when paired with corresponding changes in bottom-water oxygenation (Song et al., 2013; Meyer et al., 2011) (Fig. 8C). 5.4. Biotic responses to local environmental changes The temporal evolution of Delaware Basin water has significant implications for the response of marine fauna to environmental changes such as sea level, salinity, and oxygenation. We begin by comparing the timing of redox and salinity changes to local records of turnover in fusulineacean foraminifera (Fig. 9) (Wilde et al., 1999). Garber et al. (1989) previously implicated salinity change as the main driver for replacing large Polydiexodina fusulinids with a rapid succession of small, endemic species. Our framework supports this hypothesis; the loss of Polydiexodina correlates with an episode of elevated salinity in the mid-Capitanian. Additionally, elevated salinity at the Wordian/Capitanian boundary coincides with the loss of the fusulinid Parafusulina. In contrast, open marine conditions in the Seven Rivers and Yates platforms allowed for a ∼2 myr period of stability among fusulinid species. Sea level also plays an important role in Delaware Basin brachiopod paleoecology. Olszewski and Erwin (2009) differentiated four major groups of Permian brachiopod communities, which are separated by sea level lowstands (unconformities). The species compositions are distinct between these four groups, indicating that sea level fall was coincident with community re-organization (Olszewski and Erwin, 2009). Older brachiopod groups (Kungurian to Wordian-age) exhibit similar biogeographic and environmental gradients; however, the Capitanian-aged group lacks both generic recurrence (i.e., traits do not persist through time) as well as distinct internal gradients (Olszewski and Erwin, 2009). An ecological breakdown and lack of generic recurrence within Capitanian brachiopod subgroups coincides with the shifting geochemical stresses associated with basin restriction, stratification, and dysoxia identified here (Fig. 9). Furthermore, when this brachiopod data was analyzed at a finer temporal scale, Fall and Olszewski (2010) documented significant changes in paleocommunity composition cor-

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.9 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

9

Fig. 8. Processes that govern the distribution of carbon isotopes near carbonate platforms. A) Global changes in the carbon cycle should produce synchronous changes in both shallow- and deep-water carbonates. B) If there is a vertical gradient in carbon isotopes, changes in downslope transport through time will affect the mean and variance of carbon isotopes in deep water sections, but not shallow water sections. Small offsets between shallow- and deep-water settings correspond to maximum input from the platform top (times T1 and T3) C) If vertical environmental gradients change through time, geochemical signals may manifest differently in shallow versus deep-water sections. Distinguishing between these endmember processes requires additional sedimentological or geochemical data.

Fig. 9. Correlations between environmental and biotic changes in the Delaware Basin. Fusulinid zones are from Wilde et al. (1999), conodont zones (grey text) are from Lambert et al. (2002), and brachiopod communities are from Olszewski and Erwin (2009). Word. = Wordian, Wuch. = Wuchiapingian, J. = Jinogondolella.

responding to the large sequence boundary near the top of the Yates Fm. The authors explain community change as a product of reduced shelfal habitats and decreased connection to outside species pools; however, these shifts correlate beautifully to the onset of salinity stratification and dysoxic bottom-water conditions. The presence of chemical stresses also explain why some sea-level fall events are associated with turnover and not others.

5.5. Implications for the Guadalupian extinction in North America The integrated stratigraphic, paleontological, and geochemical framework for the Delaware Basin lend insight into larger-scale extinction patterns. Early studies of the Guadalupian extinction favored low Capitanian sea level as the extinction driver (Hallam and Wignall, 1999). This mechanism is especially attractive in north-

JID:EPSL

AID:115876 /SCO

10

[m5G; v1.261; Prn:22/10/2019; 12:39] P.10 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

west Pangea, where sea level fall would expose large areas of shallow epeiric seaways (Miller and Foote, 2009). Nevertheless, Clapham and Payne (2011) showed that the extinction was physiologically selective, which usually indicates respiratory stresses such as hypoxia (low pO2 ) or hypercapnia (high pCO2 ). In the open ocean, extinction selectivity would be an unexpected result if the sole trigger is shallow-water habitat loss during sea level fall. In contrast, sea level fall could be physiologically selective in isolated basins because it triggers drastic chemical changes in local water masses. The unusually widespread extent of Permian evaporites in both shallow- and deep-water settings (Fig. 1A) is strong evidence that chemical stresses were common during this time (Warren, 2010). Overall, the link between sea level change and chemical evolution of local water masses explains key observations about the Guadalupian extinction: increased severity in restricted settings (Miller and Foote, 2009), physiological selectivity (Clapham and Payne, 2011), and evidence for a limited global disruption of the carbon cycle (Jost et al., 2014). The preceding analysis is also compatible with more complex, multi-causal extinction mechanisms. While a full examination of the Capitanian extinction is beyond the scope of this work, we suggest several ways in which data from restricted basins might be used constrain other extinction mechanisms. For example, Emeishan volcanism may have set off a series of environmental changes resulting in hypoxia and hypercapnia (Bond et al., 2010; Clapham and Payne, 2011). Comparing the extinction patterns of epicratonic settings versus open ocean settings provides additional constraints. Warm, saline water holds less dissolved gasses than normal marine water (section 5.2) and shallow basins with smaller volumes of dissolved gasses may be more vulnerable to atmospheric change than larger oceanic water masses. In contrast, kill mechanisms such as expansion of the oceanic deep-water oxygen minimum zone may be less effective in basins that are already isolated from the global ocean. This is an intriguing possibility because most extinctions aside from the Capitanian are more pronounced in the open ocean (Miller and Foote, 2009). The difference in carbonate geochemistry between isolated basins and the open ocean can provide significant insight into how environmental shifts drive biotic crises, especially during the Paleozoic and Cretaceous when these epicratonic basins were common. 6. Conclusions Carbonate geochemistry from the Delaware Basin records a sixmillion year history of local water mass evolution during basin closure. During the first half of the Capitanian Stage, the basin periodically developed elevated salinity during prolonged sea level lowstands. During the latter half of the Capitanian, a dysoxic conditions appeared as circulation decreased. As a result, carbonate δ 13 C patterns deviate from globally averaged records due to local processes such as stratification and downslope transport. Local processes must be considered before apparent δ 13 C excursions are attributed to global carbon cycle changes such as volcanism or increased burial of organic carbon. Local evolution of Delaware Basin water corresponds strongly to periods of biotic turnover and paleoecological reorganization because sea level changes triggered corresponding changes in seawater chemistry. The close association of physical and chemical processes explains why an extinction associated with low sea level would be both physiologically selective and more severe in restricted basins. The Delaware Basin offers an instructive case study of nuanced local and global paleoceanographic changes; some global signals (e.g., sea level) may be transformed into local chemical signals (e.g., restriction) in epicontinental settings. Similar analysis of stratigraphic, geochemical, and faunal patterns in other re-

stricted settings may help constrain poorly-understood Paleozoic extinctions and biotic crises. Acknowledgements We thank Nate Miller for assistance with sample preparation, analytical methods, and running samples as well as Nick Ettinger, Julia Clarke, Hima Hassenruck-Gudipati, Evan Ramos, Scott Tinker, and the Kerans lab group for discussion and reviews of early manuscript drafts. Jonena Hearst and Rod Horrocks from the National Parks Service are also graciously acknowledged for assistance and permits for sample collection in Guadalupe Mountains National Park (GUMO-2018-SCI-0014) and Carlsbad Caverns National Park (CAVE-2018-SCI-0006). Funding to BS and CK was provided by the Reservoir Characterization Research Lab, a group of industry affiliates through the Bureau of Economic Geology. Additional funding to BS was provided through the Geological Society of America (GSA) and the Society for Sedimentary Geology (SEPM) student research grant programs. Appendix A. Supplementary material Supplementary material related to this article can be found online at https://doi.org/10.1016/j.epsl.2019.115876. References Algeo, T.J., Maynard, J.B., 2004. Trace-element behavior and redox facies in core shales of Upper Pennsylvanian Kansas-type cyclothems. Chem. Geol. 206, 289–318. Babcock, L.C., 1977. Life in the Delaware basin: the paleoecology of the Lamar Limestone. In: Hileman, M.E., Mazzullo, S.J. (Eds.), Upper Guadalupian Facies Permian Reef Complex, Guadalupe Mountains, New Mexico and West Texas. Society of Economic Paleontologists and Mineralogists, Permian Basin Section Publication 77-16. Bambach, R.K., Knoll, A.H., Wang, S.C., 2004. Origination, extinction, and mass depletions of marine diversity. Paleobiology 30, 522–542. Berelson, W.M., 1991. The flushing of two deep-sea basins, southern California borderland. Limnol. Oceanogr. 36, 1150–1166. Bond, D., Wignall, P., Wang, W., Izon, G., Jiang, H.-S., Lai, X.-L., Sun, Y.-D., Newton, R., Shao, L.-Y., Védrine, S., 2010. The mid-Capitanian (Middle Permian) mass extinction and carbon isotope record of South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 292, 282–294. Brown, A., Loucks, R., 1993. Influence of sediment type and depositional processes on stratal patterns in the Permian basin-margin Lamar limestone, McKittrick Canyon, Texas. In: Loucks, R.G., Sarg, J.F. (Eds.), Carbonate Sequence Stratigraphy: Recent Developments and Applications. In: American Association of Petroleum Geologists Memoir, vol. 57, pp. 435–474. Calvert, S., Pedersen, T., 1993. Geochemistry of recent oxic and anoxic marine sediments: implications for the geological record. Mar. Geol. 113, 67–88. Chafetz, H.S., Wu, Z., Lapen, T.J., Milliken, K.L., 2008. Geochemistry of preserved Permian aragonitic cements in the tepees of the Guadalupe Mountains, West Texas and New Mexico, USA. J. Sediment. Res. 78, 187–198. Clapham, M.E., Payne, J.L., 2011. Acidification, anoxia, and extinction: a multiple logistic regression analysis of extinction selectivity during the Middle and Late Permian. Geology 39, 1059–1062. Clapham, M.E., Shen, S., Bottjer, D.J., 2009. The double mass extinction revisited: reassessing the severity, selectivity, and causes of the end-Guadalupian biotic crisis (Late Permian). Paleobiology 35, 32–50. Debelius, B., Gómez-Parra, A., Forja, J., 2009. Oxygen solubility in evaporated seawater as a function of temperature and salinity. Hydrobiologia 632, 157–165. Derry, L.A., 2010. A burial diagenesis origin for the Ediacaran Shuram–Wonoka carbon isotope anomaly. Earth Planet. Sci. Lett. 294, 152–162. Etemad-Saeed, N., Hosseini-Barzi, M., Adabi, M.H., Miller, N.R., Sadeghi, A., Houshmandzadeh, A., Stockli, D.F., 2016. Evidence for ca. 560 Ma Ediacaran glaciation in the Kahar formation, central Alborz Mountains, northern Iran. Gondwana Res. 31, 164–183. Fall, L.M., Olszewski, T.D., 2010. Environmental disruptions influence taxonomic composition of brachiopod paleocommunities in the Middle Permian Bell Canyon Formation (Delaware Basin, west Texas). Palaios 25, 247–259. Fry, B., Jannasch, H.W., Molyneaux, S.J., Wirsen, C.O., Muramoto, J.A., King, S., 1991. Stable isotope studies of the carbon, nitrogen and sulfur cycles in the Black Sea and the Cariaco Trench. Deep-Sea Res. 38, S1003–S1019.

JID:EPSL

AID:115876 /SCO

[m5G; v1.261; Prn:22/10/2019; 12:39] P.11 (1-11)

B.P. Smith et al. / Earth and Planetary Science Letters ••• (••••) ••••••

Garber, R.A., Grover, G.A., Harris, P.M., 1989. Geology of the Capitan Shelf Margin— subsurface data from the Northern Delaware Basin. In: Harris, P.M., Grover, G.A. (Eds.), Subsurface and Outcrop Examination of the Capitan Shelf Margin. In: SEPM Core Workshop, vol. 13, pp. 3–269. Goldhammer, R., 1997. Compaction and decompaction algorithms for sedimentary carbonates. J. Sediment. Res. 67, 36–56. Grasby, S.E., Beauchamp, B., Bond, D.P., Wignall, P.B., Sanei, H., 2016. Mercury anomalies associated with three extinction events (Capitanian crisis, latest Permian extinction and the Smithian/Spathian extinction) in NW Pangea. Geol. Mag. 153, 285–297. Hallam, A., Wignall, P., 1999. Mass extinctions and sea-level changes. Earth-Sci. Rev. 48, 217–250. Hampton, B.D., 1989. Carbonate sedimentology of the Manzanita Member of the Cherry Canyon Formation. In: Harris, P.M., Grover, G.S. (Eds.), Subsurface and Outcrop Examination of the Capitan Shelf Margin, Northern Delaware Basin. In: Society of Economic Paleontologists and Mineralogists Core Workshop, vol. 13, pp. 431–439. Holmden, C., Creaser, R., Muehlenbachs, K., Leslie, S., Bergstrom, S., 1998. Isotopic evidence for geochemical decoupling between ancient epeiric seas and bordering oceans: implications for secular curves. Geology 26, 567–570. Hood, A., Wallace, M.W., 2015. Extreme ocean anoxia during the Late Cryogenian recorded in reefal carbonates of Southern Australia. Precambrian Res. 261, 96–111. Immenhauser, A., Holmden, C., Patterson, W.P., 2008. Interpreting the carbon-isotope record of ancient shallow epeiric seas: lessons from the recent. In: Holmden, C., Pratt, B.R. (Eds.), Dynamics of Epeiric Seas: Sedimentological, Paleontological and Geochemical Perspectives. In: Geological Association of Canada Special Publication, vol. 48, pp. 137–174. Immenhauser, A., Kenter, J.A., Ganssen, G., Bahamonde, J.R., Van Vliet, A., Saher, M.H., 2002. Origin and significance of isotope shifts in Pennsylvanian carbonates (Asturias, NW Spain). J. Sediment. Res. 72, 82–94. Isozaki, Y., Kawahata, H., Ota, A., 2007. A unique carbon isotope record across the Guadalupian–Lopingian (Middle–Upper Permian) boundary in mid-oceanic paleo-atoll carbonates: the high-productivity “Kamura event” and its collapse in Panthalassa. Glob. Planet. Change 55, 21–38. Jacobsen, S.B., Kaufman, A.J., 1999. The Sr, C and O isotopic evolution of Neoproterozoic seawater. Chem. Geol. 161, 37–57. Jin, Y., Noble, P.J., Poulson, S.R., 2012. Paleoenvironmental and paleoecological implications of Permian (Guadalupian) radiolarian and geochemical variations in the Lamar Limestone, Delaware Basin, West Texas (USA). Palaeogeogr. Palaeoclimatol. Palaeoecol. 346, 37–53. Jost, A.B., Mundil, R., He, B., Brown, S.T., Altiner, D., Sun, Y., DePaolo, D.J., Payne, J.L., 2014. Constraining the cause of the end-Guadalupian extinction with coupled records of carbon and calcium isotopes. Earth Planet. Sci. Lett. 396, 201–212. Kerans, C., Kempter, K., 2002. Hierarchical stratigraphic analysis of a carbonate platform Permian of the Guadalupe Mountains. In: American Association of Petroleum Geologists/Datapages Discovery Series, vol. 5, CD-ROM. Korte, C., Jasper, T., Kozur, H.W., Veizer, J., 2005. δ 18 O and δ 13 C of Permian brachiopods: a record of seawater evolution and continental glaciation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 224, 333–351. Kosa, E., Hunt, D.W., 2006. Heterogeneity in fill and properties of karst-modified syndepositional faults and fractures: upper Permian Capitan Platform, New Mexico, USA. J. Sediment. Res. 76, 131–151. Kump, L.R., Arthur, M.A., 1999. Interpreting carbon-isotope excursions: carbonates and organic matter. Chem. Geol. 161, 181–198. Lambert, L.L., Wardlaw, B.R., Nestell, M.K., Nestell, G.P., 2002. Latest Guadalupian (middle Permian) conodonts and foraminifers from West Texas. Micropaleontology 48, 343–364. Lau, K.V., Maher, K., Altiner, D., Kelley, B.M., Kump, L.R., Lehrmann, D.J., Silva-Tamayo, J.C., Weaver, K.L., Yu, M., Payne, J.L., 2016. Marine anoxia and delayed Earth system recovery after the end-Permian extinction. Proc. Natl. Acad. Sci. USA 113, 2360–2365. Lau, K.V., Macdonald, F.A., Maher, K., Payne, J.L., 2017. Uranium isotope evidence for temporary ocean oxygenation in the aftermath of the Sturtian Snowball Earth. Earth Planet. Sci. Lett. 458, 282–292. Melim, L.A., Scholle, P.A., 1999. Diagenesis of the Capitan Formation forereef facies (Permian, west Texas and New Mexico). In: Saller, A.H. (Ed.), Geologic Framework of the Capitan Reef. In: Society of Economic Paleontologists and Mineralogists Special Publication, vol. 65, pp. 193–210. Melim, L.A., Scholle, P.A., 2002. Dolomitization of the Capitan Formation forereef facies (Permian, west Texas and New Mexico): seepage reflux revisited. Sedimentology 49, 1207–1227.

11

Melim, L.A., Swart, P.K., Maliva, R.G., 2001. Meteoric and marine-burial diagenesis in the subsurface of Great Bahama Bank. In: Ginsburg, R.N. (Ed.), Subsurface Geology of a Prograding Carbonate Platform Margin, Great Bahama Bank: Results of the Bahamas Drilling Project. In: SEPM (Soc. Sediment. Geol.) Spec. Publ., vol. 70, pp. 137–161. Metzger, G.J., Fike, D.A., 2013. Techniques for assessing spatial heterogeneity of carbonate δ 13 C values: implications for craton-wide isotope gradients. Sedimentology 60, 1405–1431. Meyer, K., Yu, M., Jost, A., Kelley, B., Payne, J., 2011. δ 13 C evidence that high primary productivity delayed recovery from end-Permian mass extinction. Earth Planet. Sci. Lett. 302, 378–384. Miller, A.I., Foote, M., 2009. Epicontinental seas versus open-ocean settings: the kinetics of mass extinction and origination. Science 326, 1106–1109. Nestell, G.P., Nestell, M.K., 2006. Middle Permian (Late Guadalupian) foraminifers from Dark Canyon, Guadalupe Mountains, New Mexico. Mar. Micropaleontol. 52, 1–50. Olszewski, T.D., Erwin, D.H., 2009. Change and stability in Permian brachiopod communities from western Texas. Palaios 24, 27–40. Payne, J.L., Kump, L.R., 2007. Evidence for recurrent Early Triassic massive volcanism from quantitative interpretation of carbon isotope fluctuations. Earth Planet. Sci. Lett. 256, 264–277. R Core Team, 2017. R: A Language and Environment for Statistical Computing. R Foundation for Statistical Computing, Vienna, Austria. http://www.R-project. org/. Romanek, C.S., Grossman, E.L., Morse, J.W., 1992. Carbon isotopic fractionation in synthetic aragonite and calcite: effects of temperature and precipitation rate. Geochim. Cosmochim. Acta 56, 419–430. Saltzman, M.R., 2003. Late Paleozoic ice age: oceanic gateway or pCO2 ? Geology 31, 151–154. Schroder, S., Grotzinger, J., 2007. Evidence for anoxia at the Ediacaran–Cambrian boundary: the record of redox-sensitive trace elements and rare earth elements in Oman. J. Geol. Soc. 164, 175–187. Scotese, C., Langford, R., 1995. Pangea and the paleogeography of the Permian. In: The Permian of Northern Pangea. Springer, pp. 3–19. Shen, S., Shi, G.R., 1996. Diversity and extinction patterns of Permian Brachiopoda of South China. Hist. Biol. 12, 93–110. Song, H., Tong, J., Algeo, T.J., Horacek, M., Qiu, H., Song, H., Tian, L., Chen, Z.-Q., 2013. Large vertical δ 13 C DIC gradients in Early Triassic seas of the South China craton: implications for oceanographic changes related to Siberian Traps volcanism. Glob. Planet. Change 105, 7–20. Stigall, A.L., 2010. Using GIS to assess the biogeographic impact of species invasions on native brachiopods during the Richmondian Invasion in the TypeCincinnatian (Late Ordovician, Cincinnati region). Palaeontol. Electronica 13, 5A. Stigall, A.L., Bauer, J.E., Lam, A.R., Wright, D.F., 2017. Biotic immigration events, speciation, and the accumulation of biodiversity in the fossil record. Glob. Planet. Change 148, 242–257. Swart, P.K., Eberli, G., 2005. The nature of the δ 13 C of periplatform sediments: implications for stratigraphy and the global carbon cycle. Sediment. Geol. 175, 115–129. Tinker, S.W., 1998. Shelf-to-basin facies distributions and sequence stratigraphy of a Steep-Rimmed carbonate margin: Capitan depositional system, McKittrick Canyon, New Mexico and Texas. J. Sediment. Res. 68, 1146–1174. Tribovillard, N., Algeo, T.J., Lyons, T., Riboulleau, A., 2006. Trace metals as paleoredox and paleoproductivity proxies: an update. Chem. Geol. 232, 12–32. Warren, J.K., 2010. Evaporites through time: tectonic, climatic and eustatic controls in marine and nonmarine deposits. Earth-Sci. Rev. 98, 217–268. Wilde, G.L., Rudine, S.F., Lambert, L.L., 1999. Formal designation: Reef Trail Member, Bell Canyon Formation, and its significance for recognition of the GuadalupianLopingian boundary. In: Saller, A.H. (Ed.), Geologic Framework of the Capitan Reef. In: Society of Economic Paleontologists and Mineralogists Special Publication, vol. 65, pp. 193–210. Yang, K.M., Dorobek, S.L., 1995. The Permian basin of west Texas and New Mexico: tectonic history of a “composite” foreland basin and its effects on stratigraphic development. In: Stratigraphic Evolution of Foreland Basins. In: SEPM Special Publication, vol. 52, pp. 149–174. Zhou, M.-F., Malpas, J., Song, X.-Y., Robinson, P.T., Sun, M., Kennedy, A.K., Lesher, C.M., Keays, R.R., 2002. A temporal link between the Emeishan large igneous province (SW China) and the end-Guadalupian mass extinction. Earth Planet. Sci. Lett. 196, 113–122.