Carbon and oxygen isotope stratigraphy of the Lower ... · greenhouse conditions of the early Paleozoic and the late Paleozoic ice ages. The middle Paleozoic greenhouse-icehouse transition

For permission to copy, contact [email protected] 2002 Geological Society of America96

GSA Bulletin; January 2002; v. 114; no. 1; p. 96–108; 10 figures; Data Repository item 2002015.

Carbon and oxygen isotope stratigraphy of the Lower Mississippian(Kinderhookian–lower Osagean), western United States:

Implications for seawater chemistry and glaciation

Matthew R. Saltzman*Department of Geological Sciences, Ohio State University, Columbus, Ohio 43210, USA

ABSTRACT

A positive carbon isotope (d13C) excur-sion has been recognized in upper Kinder-hookian and early Osagean carbonates inthree sections in southeast Idaho and Ne-vada. The oldest d13C peak (17‰) is datedto the isosticha conodont zone, and a youn-ger peak occurs in the typicus Zone. Theshifts are recorded in a range of carbonatelithofacies representing various waterdepths along the shelf. Lithofacies sampledfor d13C and d18O at the Samaria Mountainsection in southeast Idaho record the shal-lowest-water conditions, indicated by cross-bedded skeletal and peloidal grainstones.The deepest water conditions are present inthe Pahranagat Range section in easternNevada, which consists mainly of biotur-bated lime mudstone and skeletal wacke-stone. The d13C values from these widelyseparated sedimentary basins show a con-sistent trend that correlates with Early Mis-sissippian curves generated from brachio-pod calcite in western Europe and theMidcontinent of North America, as well asdolomites in Utah and Wyoming. d18O val-ues become more positive up section, gen-erally paralleling the positive trend in d13Cduring the late Kinderhookian. No subaer-ial exposure surfaces are recognized in thesections examined in southeast Idaho andNevada, and at least the d13C trends are in-terpreted as primary seawater fluctuations.Sea-level changes occurred near the begin-ning of the late Kinderhookian d13C shift(early to middle parts of the isosticha Zone)and within the peak of the d13C excursion(Kinderhookian-Osagean boundary), al-though tectonic changes associated with the

*E-mail: [email protected].

Antler orogeny have likely modified the eu-static signature.

Keywords: carbon-13, conodont, glaciation,Kinderhookian, Mississippian, oxygen-18.

INTRODUCTION

The Early Mississippian marked a transi-tional period in Earth history between thegreenhouse conditions of the early Paleozoicand the late Paleozoic ice ages. The middlePaleozoic greenhouse-icehouse transition hastraditionally been viewed in terms of progres-sive cooling and ice buildup throughout theCarboniferous, linked ultimately to changes incarbon cycling associated with the rise of vas-cular land plants (Berner, 1990) and to chang-es in the positions of the continents with re-spect to the poles (Crowell, 1995). Indeed, theEarly Mississippian was characterized bysmall-scale buildup of ice in South America(Hunicken et al., 1986; Garzanti and Sciun-nach, 1997) that preceded widespread ice-sheet development on Gondwana and cyclo-themic deposition in Euramerica by tens ofmillions of years (Veevers and Powell, 1987).However, an emerging d13C and d18O recordhints at a relatively abrupt oceanographicevent in earliest Carboniferous time (Mii etal., 1999; Bruckschen et al., 1999; Saltzmanet al., 2000b). These studies represent a fur-ther refinement of the Devonian-Carbonifer-ous transition that has been recognized by pre-vious investigators (Veizer et al., 1986; Poppet al., 1986; Lohmann and Walker, 1989; Car-penter and Lohmann, 1997) and that may havehad a profound impact on the overall pace andseverity of the middle Paleozoic greenhouse-icehouse transition.

A globally significant oceanographic eventduring the Kinderhookian Provincial Series(middle Tournaisian) is signaled by a large,

positive d13C shift that reaches a maximumvalue of .17‰, which is among the highestpeaks known in the Phanerozoic (Veizer et al.,1999). The late Kinderhookian d13C shift hasbeen observed by four independent researchgroups working with a range of carbonatecomponents (brachiopod calcite, micrite, re-placive dolomite) in both European and NorthAmerican sedimentary basins (Fig. 1) (Budaiet al., 1987; Bruckschen and Veizer, 1997; Miiet al., 1999, 2001; Saltzman et al., 2000b).The possibility that this positive d13C shiftmarks the onset of the Carboniferous glacia-tion was suggested by Mii et al. (1999) on thebasis of a positive shift in d18O values in bra-chiopod calcite that paralleled the d13C shiftin North America and Europe. The Kinder-hookian d18O shift (up to 13‰ in the Mid-continent region of North America) has beeninterpreted to reflect a combination of tem-perature and ice-volume effects that occurredin response to a period of enhanced organicmatter burial (high d13C) and decreased at-mospheric carbon dioxide levels (Bruckschenand Veizer, 1997; Bruckschen et al., 1999; Miiet al., 1999, 2001). However, the precise mag-nitude and relative synchroneity of the d13Cand d18O shifts in different parts of the worldare known in only limited detail, and thus acausal link to the full-blown icehouse statethat was to follow remains uncertain. In fact,one of the major hurdles in the acceptance ofparallel d13C and d18O trends as representativeof global changes during any time period isthe independent evidence of their correlationusing biostratigraphy. The purpose of this pa-per is to provide evidence for substantial shiftsin d13C and d18O values during the late Kin-derhookian in three sections in North Americathat can be well dated by using the establishedstandard conodont zonation (Sandberg et al.,1978; Lane et al., 1980) and to discuss theirprimary versus diagenetic nature.

Geological Society of America Bulletin, January 2002 97

CARBON AND OXYGEN ISOTOPE STRATIGRAPHY OF THE LOWER MISSISSIPPIAN

Figure 1. Mississippian paleogeography for the Kinderhookian and lower Osagean Pro-vincial Series (Tournaisian Stage) after Witzke (1990). Filled circles represent localities inNorth America (Nevada, Utah, Idaho, Wyoming, and Iowa) and western Europe where alate Kinderhookian d13C excursion has been detected (Budai et al., 1987; Mii et al., 1999;Bruckschen et al., 1999). Land areas are shaded.

Figure 2. Locality map showing the six North American localities discussed in text. Sam-pling for d13C at each locality has revealed a large, positive excursion during the lateKinderhookian. Sections sampled for this study include PR—Pahranagat Range, AC—Arrow Canyon Range, SA—Samaria Mountain. Sections sampled by previous workersinclude SR—Salt River Range, GC—Gilmore City, CM—Crawford Mountains (Budai etal., 1987; Mii et al., 1999).

GEOLOGIC BACKGROUND

Study Area and Sample Collection

Carbonate samples for d13C and d18O comefrom three localities spanning the (Lower Mis-sissippian) Kinderhookian Provincial Seriesand lower part of the Osagean Provincial Se-ries in the western United States (Figs. 1, 2).The limestone successions sampled includethe Joana Limestone and Limestone X in thePahranagat Range, east-central Nevada (Ala-mo section of Singler, 1987); the Crystal Pass,Dawn, and Anchor Limestones in the ArrowCanyon Range, southeast Nevada (HiddenValley section of Brenckle, 1973); and theHenderson Canyon Formation at SamariaMountain in southeast Idaho (Gardner Canyonsection of Chen and Webster, 1994) (Fig. 3).These limestone units were deposited inroughly north-trending facies belts that corre-spond generally to a distal foreland basin bor-dering the Antler orogenic highlands (Pahran-agat Range), a forebulge setting east of theforeland (Samaria Mountain), and a shelf-mar-gin facies (Arrow Canyon Range) (Fig. 4;Poole and Sandberg, 1991; Chen and Webster,1994; Giles, 1996). The foreland basin andforebulge were bordered on the west by aflysch trough and adjacent highlands (Fig. 4),all of that formed in response to thrust loadingand flexural subsidence during the contrac-tional Late Devonian to Early MississippianAntler orogeny (Giles, 1996).

Biostratigraphic Framework

The d13C curves in this study have beencorrelated by using the standard Siphonodella(Sandberg et al., 1978) and post-Siphonodellaconodont zonations (Lane et al., 1980) for the(Lower Mississippian) Kinderhookian and Os-agean Provincial Series (Fig. 3). Most of thestudy interval falls within the uppermost Kin-derhookian Siphonodella isosticha–late cren-ulata Zone (hereafter referred to as the isos-ticha Zone) and earliest Osagean Gnathodustypicus Zone (this zone was divided into earlyand late parts by Lane et al., 1980, but is here-after referred to collectively as the typicusZone). The base of the isosticha Zone is de-fined by the first occurrence of the earliestgnathodiid species Gnathodus delicatus(Sandberg et al., 1978). The base of the typi-cus Zone, and the boundary between the Kin-derhookian and Osagean Provincial Series(Fig. 3), has been placed at the first occurrenceof Gnathodus typicus morphotype 2, whichevolved from G. delicatus (Lane et al., 1980).In practice, the Kinderhookian-Osagean

boundary has also been placed just above thelast occurrence of Siphonodella or at the firstoccurrence of Polygnathus communis carina(Sandberg et al., 1978; Thompson and Fel-lows, 1970; Lane, 1974, 1978; Carman, 1987;Chen et al., 1994).

FACIES AND DEPOSITIONALENVIRONMENTS

The three stratigraphic successions exam-ined in southeast Idaho and Nevada are madeup of limestone lithofacies that are grouped

98 Geological Society of America Bulletin, January 2002

M.R. SALTZMAN

Figure 3. Chart showing stratigraphic nomenclature used in text based on the followingsources: Wickwire et al. (1985) and Chen et al. (1994), southeast Idaho and western Wy-oming; Singler (1987), eastern Nevada; Poole and Sandberg (1991) and Pierce and Lan-genheim (1974), southeastern Nevada; and Witzke and Bunker (1996), northern Iowa.Standard conodont zonation after Sandberg et al. (1978) and Lane et al. (1980). The studyinterval corresponds to the Siphonodella isosticha and Gnathodus typicus zones. Zonalboundaries are approximate, and not all zones are found in all areas. The S. sulcatathrough S. sandbergi Zones may be present in western Wyoming, southeastern Nevada,and northern Iowa, but this interpretation remains conjectural because of unfossiliferousintervals.

generally into shoal-water and deep-ramp fa-cies associations, indicative of water depthsabove and below normal wave base, respec-tively. In general, the Pahranagat Range sec-tion contains the largest proportion of deeper-water limestone facies, and the shallowestwater facies are most abundant in the SamariaMountain section (Figs. 5–7). Both the deep-and shallow-water lithofacies consist of nearlypure limestone; dolomite makes up ,5% ofthe sections, and clastic materials are restrict-ed to thin shaly partings that make up a triv-ial amount of exposed section. Lithofaciesare arranged generally in meter-scale carbon-ate cycles, with the exception of the encrin-itic (Waulsortian-like) packstones and grain-stones of the upper Joana Limestone in thePahranagat Range (Fig. 6). All three sectionscontain the full spectrum of carbonate litho-facies represented on the lithologic key inFigure 5, the general features of which aredescribed next.

Shoal-Water Facies

The shoal-water facies association includesskeletal and peloidal packstone and grain-stone, with locally abundant intraclasts and

ooids. Skeletal grains include fragments ofcrinoids and brachiopods with lesser amountsof bryozoans, foraminifera, corals, and mol-lusks. In the Samaria Mountain section, well-sorted skeletal grainstones exhibit low-anglecross-stratification, with less common high-angle cross-bedding indicative of the highest-energy conditions on the shelf. Poorly to mod-erately sorted packstone and grainstone faciesin the Samaria, Pahranagat, and Arrow Can-yon sections are planar to massively beddedand likely represent in situ skeletal banks(Chen and Webster, 1994). The observed fea-tures of the coarse-grained shoal-water faciesassociation are consistent with open-marinedeposition above normal wave base in amound-and-channel topographic setting.

Deep-Ramp Facies

The deep-ramp facies association includesbioturbated and laminated lime mudstone andskeletal wackestone. Horizontal burrows aremost common, with vertical burrows only lo-cally abundant. Bedding-plane occurrences ofZoophycus are abundant in parts of LimestoneX in the Pahranagat Range and indicate rela-tively deep-water conditions near the lower

limits of storm influence (Singler, 1987). Thin(20–30 cm) beds of moderately sorted pack-stone occur as the basal units of meter-scale,fining-upward sequences in the PahranagatRange and are interpreted as storm-generated,allochthonous deposits. Bedded and nodularchert are locally abundant, particularly in thelower part of the Samaria Mountain sectionand the upper part of the Arrow CanyonRange section. Sponge spicules are very abun-dant in the wackestones and lime mudstonesof the deep-ramp facies association, with less-er amounts of crinoids, brachiopods, andbryozoans. Corals are locally abundant andappear to be in growth position at severalstratigraphic levels in the Pahranagat Rangesection. Fine pyrite is found throughout thisfacies, particularly in association with the lam-inated mudstone facies in the PahranagatRange that may reflect intermittent depositionin poorly oxygenated waters between majorstorm events. The observed features of thefine-grained limestones that make up the deep-ramp facies association point to deposition inquiet water below normal wave base.

STABLE ISOTOPE STRATIGRAPHY

Methods

In order to generate relatively high-resolu-tion (meter-scale), continuous stable isotopecurves, the d13C and d18O values in this studywere derived from the full range of fine-grained to coarse-grained lithologies previous-ly described and illustrated in Figures 5–7.Micrites were preferentially drilled from freshrock surfaces (generating 0.5–1.0 mg of pow-der), although nearly a third of the samplesincluded coarser-grained facies cemented withsparry calcite. All samples were roasted undervacuum at 380 8C for 1 h to remove volatilecontaminants. Samples from the Arrow Can-yon and Samaria Mountain localities were an-alyzed at the University of Iowa and Univer-sity of Michigan stable isotope laboratories,where carbonate powders were reacted with100% phosphoric acid at 75 8C in an onlinecarbonate preparation line (Carbo-Kiel—sin-gle-sample acid bath) connected to a FinniganMat 251 or 252 mass spectrometer. All isotoperatios were corrected for 17O contribution andare reported in per mil relative to the Peedeebelemnite (PDB) isotope standard. Analyticalprecision for d13C and d18O based on duplicateanalyses and on multiple analyses of NBS19was #0.04‰.



Figure 4. Major paleogeographic features and facies belts of the western United Statesduring the Kinderhookian–early Osagean, simplified after Poole and Sandberg (1991),Chen and Webster (1994) and Giles (1996). See text for discussion. PR—PahranagatRange, AC—Arrow Canyon Range, SA—Samaria Mountain.

Arrow Canyon Range, SoutheasternNevada

The Early Mississippian d13C and d18O val-ues from the Arrow Canyon Range are pre-sented in stratigraphic profile in Figure 5 (alsosee Data Repository Table DR11). The sectionbegins with d13C values of ;0‰ near the baseof the Mississippian section and peak at 17‰in the Dawn Limestone. The monotonic riseand fall is interrupted near the peak of the ex-cursion, where values show a brief negativeshift to near 15‰ before rising again above16‰ (Fig. 5). After a return to near preex-cursion values below 12‰ near the base ofthe Anchor Limestone, values begin to riseagain. Key fossil collections that set age limitson shifts in d13C in the Arrow Canyon sectioninclude specimens of P. communis carina at;20 m above the base of the Anchor Lime-

1 GSA Data Repository item 2002015, LowerMississippian stable isotope data from Nevada andIdaho, is available on the Web at http://www.geosociety.org/pubs/ft2002.htm. Requestsmay also be sent to [email protected].

stone (Pierce and Langenheim, 1974), as wellas those collections in the upper Dawn andlower Anchor that contain conodonts diagnos-tic of the typicus Zone shown in the biostrati-graphic column of Poole and Sandberg(1991).

The d18O values from the Crystal PassLimestone in the basal part of the Arrow Can-yon section are relatively low at ,–8‰ (Fig.5). The d18O values of remainder of the sam-ples are more positive than27‰ and show ageneral trend toward higher values up sectionthrough the Dawn Limestone, with a total shiftof ;13‰ generally paralleling the positivechange in d13C. Values are largely invariantthrough most of the upper Dawn and lowerAnchor Formations, although the top foursamples in the Anchor form a notable positivespike above23‰.

Pahranagat Range, Eastern Nevada

The Early Mississippian d13C and d18Ocurves from the Pahranagat Range are pre-sented in Figure 6 (and Table DR2 [see foot-

note 1]). The d13C values are near 0‰ at thebase of the Mississippian section and rise to apeak of .17‰ near the base of Member Bof Limestone X. The d13C values then fallback below 14‰ before rising again to a sec-ond distinct peak above 16‰ within MemberC. The ages of inflections in the d13C curve inthe Pahranagat Range are defined by collec-tions containing S. isosticha throughout Mem-ber A of Limestone X and collections con-taining G. typicus M2 in the lower third ofMember C (Singler, 1987) (Fig. 6).

The lowest d18O values occur in the basalpart of the Pahranagat section and show abrief shift to values more negative than26‰before a gradual increase from25‰ tonear22‰ in the lower part of Member B ofLimestone X at about the level of the initiald13C peak above 17 (Fig. 6). Values then fallbriefly back below24‰ in the upper part ofMember B before rising again and hoveringgenerally between24‰ and22‰ for the re-mainder of the sampled section.

Samaria Mountain, Southeastern Idaho

The Early Mississippian d13C and d18Ocurves from the Samaria Mountain section arepresented in Figure 7 (and Table DR3 [seefootnote 1]). Values of d13C are near 11‰ atthe base of the Mississippian section and riseto a peak of .16‰ in the lower third of theBrush Canyon Member of the Henderson Can-yon Formation. This rise is interrupted by abrief negative shift back below 12‰ towardthe top of the Chinese Wall Member. Follow-ing peak values in the middle of the BrushCanyon Member, d13C values fall back to14‰ before rising up above 15‰ to form asecondary peak near the top of the member.A fall back to preexcursion values below 0‰occurs in the Devil Creek Member of the Hen-derson Canyon Formation (Fig. 7). The d13Cshifts at Samaria Mountain are dated by co-nodont collections containing G. delicatus andG. typicus M2, at 1 m and 45 m above thebase of the Chinese Wall and Brush CanyonMembers, respectively (Wickwire et al., 1985;Chen et al., 1994). In addition, abundant spec-imens of P. communis carina are found in theupper part of the Devil Creek Member (Chenet al., 1994).

The d18O values in the Samaria Mountainsection become generally more positive upsection, shifting from near28‰ to23‰ (Fig.7). The d18O values within the upper BrushCanyon and Devil Creek Members show briefshifts back below24‰, defined by only a fewpoints. The d18O curve at Samaria Mountain


M.R. SALTZMAN

Figure 5. Lithologic column and d13C and d18O results from the Arrow Canyon Range, southeastern Nevada (Fig. 2). Boundary betweenisosticha and typicus Zones based on diagnostic conodont collections in Pierce and Langenheim (1974) from the Anchor Limestone andon the biostratigraphic column in Poole and Sandberg (1991).

is generally noisier than the Pahranagat andArrow Canyon sections (Figs. 5, 6).

DISCUSSION

The discussion is divided into two parts.The first addresses the issue of diagenesis forboth the d13C and d18O curves presented inFigures 5–7. The second assesses the relativesynchroneity of important d13C and d18O shiftsand interprets them in the context of LowerMississippian biostratigraphy and sequencestratigraphy.

Primary Versus Diagenetic Signatures

Carbon IsotopesDetermining the primary versus secondary

nature of observed trends of stable isotopes inancient carbonate sequences is a longstandingdebate. The nature of the problem with respectto carbon isotopes was pointed out early onby Hudson (1975), who noted that Keith and

Weber’s (1964) comprehensive survey re-vealed that most ancient limestones have d13Cvalues similar to unlithified carbonate sedi-ment in modern marine settings (0 to 12‰relative to PDB), whereas recent limestoneslithified by meteoric diagenesis (e.g., Bermu-dan limestones) are often highly depleted in13C (their d13C values go as low as28‰).Hudson (1975) concluded that these ancientlimestones could not have been cemented bythe same processes that occur in modern, near-surface freshwater diagenetic settings, such asBermuda, which involve the addition of extra-neous CO2 to the system, and suggested alter-natively that lithification may have taken placeduring burial diagenesis and was effectively aclosed system with respect to carbon. In the1980s and 1990s, quantitative studies of car-bonate diagenesis and water-rock interactionadded to the debate by demonstrating thatd13C values are likely to be preserved over awide range of water:rock ratios and degrees ofdiagenetic alteration, including some involv-

ing near-surface diagenesis, because diagenet-ic fluids acquire most of their carbon directlyfrom the dissolution of metastable carbonates(Magaritz, 1983; Lohmann, 1988; Banner andHanson, 1990). Nonetheless, there are docu-mented cases in which the d13C values of an-cient limestones have been shifted towardvery low values in close proximity (usuallywithin ,10 m) to known exposure surfaces,likely reflecting the addition of oxidized or-ganic carbon during meteoric diagenesis (Al-lan and Matthews, 1982; Algeo et al., 1992.Thus, it is important to place the d13C shiftsin this paper in a sequence stratigraphic frame-work that can be used to evaluate the presenceor absence of exposure surfaces.

The limestone successions sampled for sta-ble isotopes (Figs. 5–7) consist of shoal-waterand deep-ramp lithofacies that are interpretedto reflect deposition in entirely subtidal (per-manently submerged) environments. Even theshallowest water facies in the Samaria Moun-tain section (e.g., cross-bedded skeletal and



Figure 6. Lithologic column and d13C and d18O results from the Pahranagat Range, south-eastern Nevada (Fig. 2). Boundary between isosticha and typicus Zones based on diagnosticconodont collections from Limestone X documented in Singler (1987). See Figure 5 forstratigraphic legend.

oolitic grainstones) do not preserve evidenceof subaerial exposure, such as pedogenic fea-tures (sandy, oxidization crusts), karsticweathering (solution-enlarged joints and pits),fenestrae, mud cracks, or evaporite deposition

(Allan and Matthews, 1982; Tucker andWright, 1990; Algeo et al., 1992). The mostprominent stratigraphic surfaces in the sec-tions sampled, at the tops of the Joana Lime-stone in the Pahranagat Range and Crystal

Pass Formation in the Arrow Canyon Range,are best described as marine flooding surfacesand locally show evidence for submarinehardground formation and condensation (co-nodont-rich marker horizon of Giles, 1996).These surfaces coincide with minor d13C shifts(,0.5‰) or no shift at all, inconsistent withprolonged exposure. The most distinct strati-graphic interval of lowered d13C values occursnear the top of Member B of Limestone X inthe Pahranagat Range (Fig. 6) within rocks ofthe deep-ramp facies associations that arefound throughout the section and are indica-tive of deposition below normal wave base.Similarly, the relatively low d13C values nearthe top of the Dawn Limestone in the ArrowCanyon Range occur within wackestone andpackstone lithologies that show no evidenceof exposure. More generally, the relativelysmooth anatomy of the trends of decreasingd13C values over tens of meters of section, asshown in Figures 5–7, do not appear consis-tent with episodes of subaerial exposure,which are typically marked by short-livedd13C shifts (on the order of meters leading upto exposure surfaces) that produce abruptbreaks in the d13C trend (Allan and Matthews,1982; Algeo et al., 1992).

An additional line of evidence that wouldseem to argue for the primary nature of d13Cshifts seen in this study is based on estimatesof the d13C value of normal Early Mississip-pian seawater. In order for the large d13C shiftsin southeast Idaho and Nevada (Figs. 5–7) tobe attributed entirely to the effects of meteoricdiagenesis, one would need to argue that nor-mal Mississippian seawater was near 17‰.This conclusion appears to be inconsistentwith published d13C values for Mississippianseawater of #14‰ (Brand, 1982; Meyers andLohmann, 1985; Popp et al., 1986; Mii et al.,1999). Values of d13C as high as 17‰ arevery rarely encountered in Phanerozoic lime-stones and occur exclusively at the peaks ofsimilar positive d13C excursions during theLate Ordovician (Kump et al., 1999; Finneyet al., 1999) and Late Silurian (Azmy et al.,1998), reminiscent of the large d13C shifts inthe Neoproterozoic (Knoll et al., 1986; Derryet al., 1992; Kaufman and Knoll, 1995; Grot-zinger et al., 1995; Hoffman et al., 1998; Ken-nedy et al., 1998).

The interpretation of d13C trends in this pa-per as primary is also consistent with the in-creasing number of published studies in thePaleozoic and Precambrian that have reportedsimilar d13C trends from widely separatedstratigraphic sequences by using a range ofcarbonate components (Gao and Land, 1991;Ripperdan et al., 1992; Joachimski and Bug-


M.R. SALTZMAN

Figure 7. Lithologic column and d13C and d18O results from Samaria Mountain, southeastern Idaho (Fig. 2). Boundary between isostichaand typicus Zones based on diagnostic conodont collections from the Henderson Canyon Formation in Wickwire et al. (1985) and Chenet al. (1994). See Figure 5 for stratigraphic legend.

gisch, 1993; Saltzman et al., 1995, 1998,2000a; Pelechaty et al., 1996; Wang et al.,1996; Kaljo et al., 1997; Patzkowsky et al.,1997; Glumac and Walker, 1998; Saylor et al.,1998; Kump et al., 1999), including marinecements (Carpenter and Lohmann, 1997) andbrachiopod calcite (Marshall et al., 1997;Azmy et al., 1998). It is particularly encour-aging that comparisons of d13C curves gener-ated from brachiopod calcite and micritesyield similar trends not only for the Missis-sippian, but for the Ordovician and Silurian aswell. For example, in the Ordovician of NorthAmerica (Finney et al., 1999; Kump et al.,1999) and the Silurian of Europe (Kaljo et al.,1997; Wigforss-Lange, 1999) and NorthAmerica (Saltzman, 2001), essentially bulk-rock analyses of micritic limestones faithfullyrecord several of the major positive d13C ex-cursions that have been recognized in brachio-pod calcite worldwide (Wenzel and Joachim-ski, 1996; Marshall et al., 1997; Bickert et al.,1997; Azmy et al., 1998).

Oxygen IsotopesIn contrast to d13C, d18O values are com-

monly not presented in chemostratigraphicstudies of ancient limestones (e.g., Ripperdan

et al., 1992; Kennedy et al., 1998), in part be-cause water-rock modeling and empiricalstudies have shown that the d18O signature canbe overprinted by even small pore volumes ofmeteoric diagenetic fluids even where the d13Csignature is preserved (Banner and Hanson,1990; Carpenter and Lohmann, 1997). Valuesof d18O that have been presented for Paleozoicmicritic or bulk-rock limestones often showno trend when plotted against major d13C ex-cursions (Saltzman et al., 1998; Kump et al.,1999), consistent with modeling predictions.However, recent work on Paleozoic brachio-pods has consistently provided evidence forsubstantial positive shifts in d18O values as-sociated with large positive d13C excursions inthe Ordovician (Brenchley et al., 1994; Mar-shall et al., 1997), Silurian (Wenzel and Joach-imski, 1996; Bickert et al., 1997; Azmy et al.,1998), and Mississippian (Bruckschen et al.,1999; Mii et al., 1999). Although the presenceof covariant d18O and d13C trends has beentraditionally associated with meteoric diagen-esis in that the more negative values representsamples stabilized in meteoric waters that hadless 18O and contained 12C from respiration ofsubaerial vegetation (Hudson, 1977; Allan andMatthews, 1982; Lohmann, 1988), Paleozoic

brachiopod workers have viewed these sametrends as indicators of global shifts in biogeo-chemical cycling. Specifically, when consid-ered in a geologic framework that includes ev-idence of glaciation, positive covariance istaken to indicate episodes of enhanced burialof 12C-enriched organic matter; such episodesresult in drawdown of atmospheric CO2 andglobal cooling (see also discussion of covari-ant stable isotope trends in Marshall, 1992;Veizer et al., 1999).

The d18O profiles in all three Lower Mis-sissippian sections in Figures 5–7 show trendstoward more positive values up section in theupper Kinderhookian, paralleling positivetrends in d13C and showing a covariant rela-tionship (Tables DR1–DR3 [see footnote 1])as already discussed. The positive shift in d18Ois best displayed in the Arrow Canyon Rangesection where values change from26‰ in thelower part of the Dawn Limestone to23‰ inthe middle part of the Dawn, whereas over thesame interval, the carbon changes from 14 to17‰. The d18O values begin to decline againin the upper part of the Dawn Limestone,again in step with d13C (Fig. 5). The trends ind18O show important similarities with trendsderived from studies of time-equivalent bra-



chiopod calcite. The d18O results in Brucksch-en et al.‘s (1999) brachiopod calcite studyshow a transition from values of ,–6‰ tomore positive values in the late Kinderhookian(Tournaisian CI to CII transition in their Fig.9), and Mii et al.‘s (1999) d18O values changefrom near25‰ (although these low valuesfrom the Glen Park Formation did not meetstrict preservation criteria) to21‰ over thesame time interval (D-CI transition in theirFigs. 6 and 8). Differences in the magnitudeof the d18O shifts during the late Kinderhook-ian in these different studies could be due tobiostratigraphic uncertainty, local seawatervariability, or species or diagenetic effects. Onthe basis of the available data, it seems likelythat at least the most negative d18O values inthe lowermost parts of the three sampled sec-tions represent signatures of burial (high-tem-perature) diagenesis. An entirely diageneticinterpretation of the remainder of the d18Otrends presented in this paper must remain anopen question pending (1) further study of thediagenetic pathways of the lithologies sam-pled and (2) comparison with published trendsgenerated from well-preserved brachiopods orcalcite cements.

In summary, it is viewed as encouragingthat the changes in d13C and d18O values re-corded in this study (Figs. 5–7) show impor-tant similarities with published curves gener-ated by using distinctly different samplingstrategies, including (1) microsampling of re-placive dolomite, crinoidal calcite, calcite ce-ment, and dolomite cement from the shelf-margin facies of the Madison Group inWyoming and Utah by Budai et al. (1987) and(2) the isolation of petrographically well-pre-served brachiopod calcite in the U.S. Midcon-tinent (Mii et al., 1999) and western Europe(Bruckschen and Veizer, 1997; Bruckschen etal., 1999). The remaining section focuses onthe correlation of d13C shifts (Fig. 8) with im-portant biostratigraphic and sequence strati-graphic boundaries.

Mississippian Biostratigraphy and d13C

Making the assumption that the large shiftsin d13C identified in this and other studies inNorth America and Europe (Fig. 9) are reflec-tive of primary seawater values, an importantquestion arises as to the relative synchroneityof the shifts in the different oceanic basins.Conodont biostratigraphy has been used wide-ly in the correlation of Lower Mississippiandepositional successions: the global standardSiphonodella (Sandberg et al., 1978) and post-Siphonodella zonations (Lane et al., 1980)subdivide the Kinderhookian and Osagean

Provincial Series, respectively. The conodontbiostratigraphy at Samaria Mountain (Fig. 7)is the most intensively studied, and the bound-aries of the isosticha and typicus zones havebeen documented by Wickwire et al. (1985)and Chen et al. (1994). The initial d13C peak(.16‰) at Samaria Mountain occurs withinthe boundaries of the isosticha Zone, markedby collections containing G. delicatus and G.typicus M2, at 1 m and 60 m above the baseof the Henderson Canyon Formation, respec-tively; and the second peak (.15‰) is withinthe typicus Zone. The minimum in d13C be-tween the peaks is poorly dated owing to asparsely fossiliferous nondiagnostic intervalthat characterizes the 15 m of section belowthe first occurrence of G. typicus M2, preclud-ing a precise location of the Kinderhookian-Osagean boundary within the Brush CanyonMember of the Henderson Canyon Formation(Fig. 7). The ages of the inflections in the d13Ccurves for the Pahranagat and Arrow Canyonsections are also consistent with peaks in theisosticha and typicus Zones (Fig. 8), as arguedsubsequently.

In the Pahranagat Range, evidence forwhich part of the d13C curve occupies the isos-ticha Zone is based on (1) six collections con-taining specimens of S. isosticha in the lower55 m of Limestone X and (2) two collectionscontaining specimens of G. typicus M2 fromhorizons at 100 and 126 m above the base ofLimestone X (Singler, 1987). The initial d13Cpeak (17‰) lies at 55–60 m above the baseof Limestone X (Fig. 6) and cannot be olderthan the isosticha Zone. Although specimensof G. delicatus, the defining species for the S.isosticha Zone (S. isosticha itself ranges intothe early crenulata Zone), were not recovered,an early crenulata Zone age for the initial d13Cpeak is precluded by two factors: (1) the gen-eral lack of conodont diversity (Sandberg etal., 1978; Ziegler and Lane, 1987) and (2) therecognition that the early crenulata Zone iseverywhere represented by a very thin interval(averaging ;3 m in thickness), which hasbeen recovered only from the basal beds ofthe Paine Member of the Lodgepole Lime-stone and from lag deposits in the underlyingCottonwood Canyon Member (Klapper, 1966;Sandberg, 1979). The younger d13C peak(.16‰) in the Member C of Limestone X isabove specimens containing G. typicus M2and thus lies within the typicus Zone (Fig. 6),as seen at Samaria Mountain (Fig. 7). In boththe Pahranagat and Samaria Mountain sec-tions, the minimum in the d13C curve betweenthe two peaks is poorly dated because of aninterval yielding only long-ranging, undi-agnostic conodont taxa.

Although conodont collections containingspecimens of Siphonodella are not knownfrom the Arrow Canyon Range, similar argu-ments based on zonal thickness and the recov-ery of important collections containing P.communis carina from the lower part of theAnchor Limestone (Pierce and Langenheim,1974) place the 17% d13C peak within theisosticha Zone (Fig. 5). This age is furthersupported by a sample containing G. delicatusnear the top of the Dawn Limestone (Goebel,1991) and is consistent with the conodont zo-nation presented in Poole and Sandberg(1991). Further study will be necessary to bet-ter locate the position of the Kinderhookian-Osagean boundary in the Arrow CanyonRange, and the zonal boundary drawn in Fig-ure 5 must be considered tentative.

Mississippian Sequence Stratigraphy andd13C

The correlation between the late Kinder-hookian d13C excursion (Fig. 8) and sea-levelchange is important in gaining a better under-standing of potential controls on Early Mis-sissippian carbon cycling, particularly in re-gards to the possible causal connectionbetween decreased atmospheric carbon diox-ide contents (and concomitant high d13C), andthe onset of ice buildup in Gondwana (Mii etal., 1999). Comparison of relative sea-levelcurves between sections in the western andMidcontinent regions of North America mayprovide one of the more reliable tests of EarlyMississippian eustasy available because theregions have been reasonably well correlatedwith each other through the use of conodontbiostratigraphy but occur in distinctly differ-ent paleotectonic settings. In the PahranagatRange in Nevada, a significant sequenceboundary occurs at the top of the Joana Lime-stone (Fig. 6) and marks a regional shift fromprogradational to retrogradational parase-quence sets (sequence 1–sequence 2 boundaryof Giles, 1996; equivalent to the Morris-Sad-lick sequence boundary of Silberling et al.,1997). This sequence boundary occurs withinthe isosticha Zone, near the beginning of therising limb of the d13C peak, and has beeninterpreted by Giles (1996) as a tectonicallyforced retrogradation resulting from flexuralsubsidence in the distal Antler foreland basin.The possibility that this boundary was en-hanced by a eustatic rise is supported by cor-relation with a significant cycle boundary inthe Midcontinent (the base of the ChouteauLimestone in Illinois, equivalent to the baseof cycle 3, Witzke and Bunker, 1996), as wellas sections in western Europe and Russia


M.R. SALTZMAN

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(Ross and Ross, 1988). This eustatic rise mayhave contributed to the onset of the d13C ex-cursion through enhanced burial of organiccarbon in deep, stratified foreland basins(Saltzman et al., 2000b).

The most prominent sequence boundary inthe Lower Mississippian occurs at the Kin-derhookian-Osagean boundary in the Midcon-tinent region (the top of the Chouteau andWassonville Limestones in southeast Iowa andthe base of cycle 4, Witzke and Bunker, 1996);this sequence boundary coincides with thed13C peak in the western United States. Thissequence boundary has also been recognizedby Lane (1978) and Ross and Ross (1988) (seeFig. 10) in time-equivalent sections in NewMexico, western Europe (Tn2-Tn3 boundary)and Russia (top Cherepetian), although it isnot consistently expressed in southeast Idahoor Nevada (Giles, 1996). A shallowing eventat the Kinderhookian-Osagean boundary isrecognized in the Pahranagat Range by a cor-al-rich zone in Limestone X (Fig. 6); however,this relative fall is not apparent at equivalentlevels in the Arrow Canyon Range or at Sa-maria Mountain. The degree of tectonic over-printing of the eustatic signature in these west-ern sections, related to changes in the rates ofsubsidence of the foreland basin or uplift atthe forebulge (Chen et al., 1994; Giles, 1996;Link et al., 1996), is difficult to assess. TheSr isotope curve of Bruckschen et al. (1995)indicates a shift to more radiogenic ratios nearthe Kinderhookian-Osagean boundary (super-imposed on a longer term decline) that couldreflect increased continental weathering dur-ing sea-level fall (Fig. 10); however, asBruckschen et al. (1995) have pointed out,tighter biostratigraphic controls will be nec-essary in attempts to evaluate the significanceof Sr isotope ratios in the context of Tournai-sian paleoceanography. Future sequence strati-graphic studies that integrate chemostratigra-phy and biostratigraphic means of correlationwill be needed to shed light on the signifi-cance of the major sequence boundary (glacio-eustatic?) at the boundary between the Kin-derhookian Provincial Series and the OsageanProvincial Series in the Midcontinent regionof the United States and its relationship to thelarge positive shift in d13C at time-equivalentlevels in the western United States.

CONCLUSIONS

Three Lower Mississippian (Kinderhookianand lower Osagean) stratigraphic sections rep-resentative of different paleotectonic and pa-leogeographic settings in the western UnitedStates show similar trends in both d13C and


M.R. SALTZMAN

Figure 10. Stratigraphic nomenclature for North America and Europe plotted against trends in d13C, d18O, Sr isotopes, and sea level.d13C curve is a composite from this study and the work of Budai et al. (1987) and Mii et al. (1999). Note d13C and d18O covariation aswell as the sea-level change near the d13C peak at the Kinderhookian-Osagean boundary. Also, note Sr shift to more radiogenic ratiosnear the Kinderhookian-Osagean boundary (superimposed on a longer term decline), which could reflect enhanced continental weath-ering during sea-level fall.

d18O values. Biostratigraphic correlation usingthe standard North American conodont zona-tion indicates synchroneity of two closelyspaced d13C peaks .17‰ in the late Kinder-hookian isosticha conodont zone and the earlyOsagean typicus Zone. The d13C shifts couldreflect meteoric diagenesis; however, the sec-tions examined in southeast Idaho and Nevadalack evidence of exposure surfaces (or alter-ation by ground water in phreatic zones). Fur-thermore, the shifts have also been recognizedin well-preserved brachiopods from westernEurope and the Midcontinent region of NorthAmerica. The cause of the d13C shift is un-certain, but may involve a significant amountof organic carbon burial in deep-marine basinsassociated with subsidence during the Antler

orogeny and eustatic rise. A significant fall insea level occurred in the Midcontinent ofNorth America, near the peak of the d13C ex-cursion in the western United States (Kinder-hookian-Osagean boundary) and could reflectglaciation. However, this sea-level event is notrecognized in western United States becausetectonic changes associated with the Antlerorogeny have likely modified the eustatic sig-nature. The primary versus secondary natureof the positive d18O shifts associated with thed13C excursion is unclear at this time and re-quires further evaluation.

ACKNOWLEDGMENTS

Detailed, thoughtful, and constructive reviewswere provided by P. Knauth, J. Humphrey, W. Mey-

ers, and P. Choquette. G. Klapper, B. Witzke, G.Ludvigson, S. Carpenter, and P. Heckel provideduseful discussion about Carboniferous and Paleo-zoic stratigraphic problems. G. Webster providedguidance to the Samaria Mountain section. A. Tay-lor, C. Kerr, and A. Koehler provided help with fieldwork in Nevada. This work was supported by Na-tional Science Foundation grant EAR-12385.

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