(barremian-albian) sections from the sierra madre ... - Instituto Longoria

Journal of Foraminiferal Research, v. 29, no. 4, p. 418–437, October 1999

THE RECORD OF GLOBAL CHANGE IN MID-CRETACEOUS (BARREMIAN-ALBIAN) SECTIONS FROM THE SIERRA MADRE, NORTHEASTERN MEXICO TIMOTHY J. BRALOWER1, EMILY COBABE2, BRADFORD CLEMENT3, WILLIAM V. SLITER4*, CHRISTOPHER L. OSBURN5, and JOSE´ LONGORIA3 ments, possibly reflecting a decrease in dilution as a result of the rise in relative sea level.

ABSTRACT

Our current understanding of mid-Cretaceous global change is largely based on investigations of pelagic sections from southern Europe and deep sea drilling sites. Much less information exists from other continents and from hemipelagic sections deposited on continental margins. This investigation seeks to broaden our understanding of mid-Cretaceous global change by focusing on the record from hemipelagic sections deposited along the continental margin of northeastern Mexico. The major goals are to compare the record, timing, and extent of the Oceanic Anoxic Events (OAEs) in Mexico and other areas, and to determine the relationship between these events and the global burial of organic material using carbon isotopes. We have investigated four sections from the Sierra Madre Oriental, integrating biostratigraphy, magnetostratigraphy and carbon isotope stratigraphy. Carbon isotopes, measured on the organic carbon (Corg ) fraction, show identical stratigraphic changes to curves from Barremian to lower Albian European and Pacific deep-sea sections. Our results add new detail to the C-isotope stratigraphy of the middle and late Albian interval. Three abrupt peaks in Corg content correlate with OAE1a (early Aptian), OAE1b (early Albian) and an event in the late Aptian Globigerinelloides algerianus Zone. All three events are marked by short-term, 0.5–3 per mil decreases in C-isotope values followed by increases of similar magnitude. The decreases may reflect changes in the type of Corg, the nature of carbon cycling, or an increase in hydrothermal activity. The increases in C-isotope values reflect widespread burial of Corg. The similar shape of the C-isotope curves in Mexico and other areas, and the response of C-isotopes to the OAEs, indicate that the late Aptian episode was extensive, and that OAE1a and OAE1b were global. The three anoxic events appear to correlate with rising relative sea level. OAE1a also corresponds to major changes in nannofossil assemblages; the well-known ‘‘nannoconid crisis’’ can be easily recognized in the Mexican sections. This event is characterized by an increase in abundance of nannofossils and foraminifera in sedi-

INTRODUCTION: MID-CRETACEOUS GLOBAL CHANGE The mid-Cretaceous is known as one of the best ancient examples of greenhouse climate (e.g., Barron and Washington, 1982; Parrish and Curtis, 1982; Berner and others, 1983). This interval was characterized by global warmth and low latitudinal temperature gradients (e.g., Douglas and Savin, 1975; Brass and others, 1982; Barron and Peterson, 1990; Huber and others, 1995; Norris and Wilson, 1998; Fassell and Bralower, 1999). The mid-Cretaceous greenhouse was coincident with a world-wide pulse in ocean crustal production (Larson, 1991, Tarduno and others, 1991; Arthur and others, 1991; Erba and Larson, 1991; Bralower and others, 1994). This, the largest volcanic episode possibly in the past 250 m.y. of Earth history, included increased rates of seafloor spreading and increased rates of formation of Large Igneous Province (LIP) oceanic plateaus, seamount chains and continental flood basalts (Hays and Pitman, 1973; Schlanger and others, 1981). The release of mantle CO2 from this enormous volcanic episode may have directly caused mid-Cretaceous greenhouse warming (Arthur and others, 1985a; Arthur and others, 1991; Larson, 1991). Mid-Cretaceous oceanic environments favored the deposition and burial of organic carbon-rich sediments, known informally as ‘‘black shales’’ (e.g., Schlanger and Jenkyns, 1976; Ryan and Cita, 1977; Arthur and Premoli Silva, 1982). Warm deep-water temperatures and sluggish circulation led to widespread dysoxia and anoxia in deep water masses (e.g., Brass and others, 1982; Wilde and Berry, 1982). Stratigraphic investigations have led to the recognition that mid-Cretaceous black shales were not deposited randomly through time, but that accumulation was concentrated in intervals known as ‘‘Oceanic Anoxic Events’’ (OAEs) (Schlanger and Jenkyns, 1976; Jenkyns, 1980; Arthur and others 1990; Bralower and others, 1993). The brief (,1 m.y.) OAE that took place at the Cenomanian/Turonian boundary (OAE2) was global in extent (e.g., Schlanger and others, 1987; Huber and others, 1999). The lengthy (;20 m.y.) Aptian-Albian OAE (OAE1) was characterized by at least four separate phases of organic-carbon accumulation (e.g., Arthur and others, 1990; Bralower and others, 1993; Erbacher and Thurow, 1997), but only the first of the four (early Aptian event [OAE1a]) is known to be global in extent (e.g., Sliter, 1989; Bralower and others, 1994). Carbon isotope records can be used to establish global organic carbon (Corg) budgets during these OAEs. The prominent positive carbon isotopic excursion that corresponds to OAE2 (e.g., Scholle and Arthur, 1980; Pratt and Threlkeld, 1984; Arthur and others, 1987) suggests that the volume of

1 Department of Geological Sciences, University of North Carolina, Chapel Hill, NC 27599-3315, USA. 2 Biogeochemistry Laboratories, Department of Geosciences, University of Massachussets, Amherst, MA 01003, USA. 3 Department of Geology, Florida International University, Miami, FL 33199, USA. 4 U. S. Geological Survey, 345 Middlefield Rd., Menlo Park, CA 94025, USA; *published posthumously. 5 Department of Earth and Environmental Sciences, Lehigh University, Bethlehem, PA 18015, USA.

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Corg buried was a sizable part of the global carbon budget. Carbon isotope records for the Aptian-Albian interval are complex (e.g., Weissert and others, 1985, 1998; Pratt and King, 1986) and their relationship with dysoxic/anoxic intervals is not fully understood. Carbon isotope values increase just above early Aptian OAE1a and remain high into the Albian. Major discrepancies exist, however, in the chronostratigraphic correlation of the C-isotope record in Mexico (Scholle and Arthur, 1980) and Europe (e.g., Weissert and Lini, 1991; Weissert and Bre´he´ret, 1991). Before short-term changes in Aptian-Albian carbon budgets can be fully explained, the age of major changes in C-isotope ratios and their relationship with Corg-burial events must be established. Oceanic environments and their microplankton inhabitants are intimately related; the evolution of microfossil groups is thought to have been profoundly affected by major oceanic events such as OAEs (e.g., Roth, 1987; Leckie, 1989; Leckie and others 1998). One of the most dramatic turnovers in the nannoplankton took place in the early Aptian around the time of OAE1a. This event involved the temporary collapse of the nannoconids, a group of nannoplankton that had dominated assemblages for the previous 20 million years, and is known as the ‘‘nannoconid crisis’’ (Coccioni and others 1992; Erba, 1994). A possible explanation for the crisis was that a mantle superplume event (e.g., Larson, 1991) directly or indirectly caused a change in the thermal or nutrient structure of oceanic surface waters (Erba, 1994). While it is suspected that the nannoconid crisis was a global event, all of the supporting data come from sections in Europe and the central Pacific Ocean that were deposited in pelagic environments. Much of our knowledge on the diverse aspects of midCretaceous global change derives from investigations of the classical Tethyan sections exposed in France and Italy (e.g., Arthur and Premoli Silva, 1982; Weissert and others 1985; Premoli Silva and others, 1989; Coccioni and others, 1992). Yet, the sedimentary record of other areas clearly holds clues to the causes of environmental and evolutionary changes. The classic paper of Scholle and Arthur (1980) that demonstrated the potential of C-isotopes in stratigraphy was largely based on the study of two sections from northeastern Mexico, but this area has received little paleoceanographic study since. This paper describes an investigation of various aspects of global change based on the study of sedimentary units in northeastern Mexico. GEOLOGICAL SETTING OF STUDIED SECTIONS TECTONICS

AND

PALEOGEOGRAPHY

The sections we investigated are exposed in ranges of the Sierra Madre Oriental fold and thrust belt of the northern Mexican Cordillera (Humphrey, 1956; Longoria, 1998) (Fig. 1). During the Barremian-Albian, northeastern Mexico was dominated by a series of interfingering shallow basins and shallow-water carbonate platforms lying on the northwest margin of a deeper seaway (e.g., Smith, 1981; Wilson, 1989). Environments of deposition varied through time as a result of tectonic processes, eustatic fluctuations and variations in detrital supply (e.g., Goldhammer and others, 1991). These sections provide a transect from pelagic (La Boca

FIGURE 1. Map of northeastern Mexico showing location of sections investigated.

Canyon) to hemipelagic (Canyon Los Chorros) deposition. Deformation of this region occurred in the Laramide orogeny (e.g., Humphrey, 1956; De Cserna, 1956; Lopez-Ramos, 1980; Tardy, 1980; Longoria, 1984; 1985; 1998), exposing the sections in elongate anticlines that form the ranges. BIOSTRATIGRAPHY

AND

LITHOSTRATIGRAPHY

The Barremian to Albian succession of northeastern Mexico contains abundant planktonic foraminifera with a wellestablished biostratigraphy (Longoria, 1974; Longoria and Gamper, 1977; Longoria, 1984; 1998). However, investigations of nannofossil stratigraphy in northeastern Mexico were carried out before most taxonomy was developed and were limited largely to the nannoconids. Trejo (1960, 1975) demonstrated that the nannoconids have significant biostratigraphic potential, at least on a regional basis. Lithostratigraphic nomenclature of the Lower Cretaceous sedimentary units of northeastern Mexico is by no means standard (e.g., Ross, 1981; Smith, 1981; Wilson and others, 1984; Longoria, 1998; Lehmann and others, in press). The lithostratigraphy of the four sections investigated is described in detail in Longoria and others (in prep.) (Fig. 2).

San Angel Limestone This unit consists of a monotonous succession of thickto massive-bedded, dark-grey to black, carbonaceous lime-

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BRALOWER, COBABE, CLEMENT, SLITER, OSBURN, AND LONGORIA

1975; Longoria and Gamper, 1977; Longoria and Monreal, 1991; Longoria, 1998).

Tamaulipas Limestone The Tamaulipas Limestone is composed predominantly of grey, medium- to thick-bedded limestone with abundant black chert layers, but contains a few mudstone and shale intervals. The unit was deposited in a basinal setting (e.g., Ross, 1981; Wilson, 1981). Planktonic foraminifera are abundant in the limestones of the Tamaulipas indicating that its age ranges from early to late Albian.

Cuesta del Cura Formation The Cuesta del Cura is composed of dark grey to black, thin- to medium-bedded limestone, with shale interbeds and chert lenses. The unit was deposited in a hemipelagic slope setting (Ice, 1981). The limestones are rich in planktonic foraminifera that provide a late Albian to early Turonian age. SECTIONS INVESTIGATED We investigated four sections in the Sierra Madre Oriental (Fig. 2). Detail on the locations of the sections is provided in Longoria and others (1998).

Santa Rosa Canyon FIGURE 2. General lithostratigraphic scheme for the BarremianCenomanian interval of the sections investigated (after Longoria, 1998).

stones and marly limestones with interbedded chert (Longoria and Davila, 1979). The San Angel was deposited in a deep-water setting and contains common planktonic foraminifers which indicate a Barremian to Aptian age (e.g., Longoria, 1998). This unit is often referred to as the lower Tamaulipas Limestone. The stratotype of this unit is exposed at Santa Rosa Canyon.

Cupido Formation This unit is a platform equivalent of the San Angel Limestone found in the west of the study area (Fig. 2). The Cupido consists of homogeneous medium- to thick-bedded light grey limestone. Sedimentary facies suggest deposition on a carbonate ramp in supratidal to lagoonal environments (e.g., Wilson, 1981; Longoria and Monreal, 1991; Lehmann and others, in press).

La Penã Formation The La Penã Formation consists of thin- to medium-bedded, dark grey and black limestone, marlstone, and shale. The unit contrasts lithologically with limestones above and below, representing an important paleoceanographic event as well as a transgressive systems tract (Goldhammer and others, 1991). This unit is rich in ammonites and planktonic foraminifera which indicate an age ranging from early Aptian to early Albian (e.g., Humphrey, 1956; Longoria, 1974,

The Santa Rosa Canyon section is exposed in the eastern front range of the Nuevo Leone Cordilliera on both sides of Highway 58 between Iturbide and Linares. The lithostratigraphy of the section has been described in detail by Ice (1981), Ross (1981), Wilson (1981), and Carslen (1998).

La Boca Canyon The section is located in the Cerro de le Silla, a range of the isolated Sierras Tamaulipecas (Longoria and others, 1998). The canyon dissects an elongate, narrow anticline that exposes the entire Mesozoic section. The lithostratigraphy of the section has been described in detail by Longoria and Davila (1979). We sampled the uppermost San Angel Limestone and La Penã Formation along the road immediately to the west of Rancho Paraiso.

Canyon Los Chorros The section is located in the Sierra de la Nieve on old Highway 57 from Matehuala to Saltillo. Exposure is continuous, but lithostratigraphy indicates that the sequence is overturned: the Cupido Formation overlies the La Penã Formation which rests on top of the Tamaulipas Limestone (see Plate 2 in Longoria and others, 1998).

Cienega del Toro The section is located in the eastern front range of the Nuevo Leone Cordilliera on a dirt road from San Pablo to Rajones near the town of Cienega del Toro (see Fig. 19 in Longoria and others [1998] for detailed location). The sampled part of the section includes the San Angel Limestone and the lower part of the La Penã Formation.

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METHODS

GEOCHEMISTRY

LITHOSTRATIGRAPHY

We collected samples for geochemical analysis every one meter in all sections and at a 0.5 meter spacing in critical intervals. Approximately 0.5 kg of sample was collected. The outer surface of field samples was removed. Samples were powdered in a ceramic shatter box and stored in an annealed glass jars. All geochemical data are archived at the World Data Center-A for Paleoclimatology.

All sections were measured and logged in detail (Figs. 3– 6). The La Boca Canyon and Cienega del Toro sections are continuously exposed, however the Santa Rosa Canyon and Canyon Los Chorros sections have structural complexities that deserve careful consideration. The sampled part of the Santa Rosa Canyon section consists of three separate segments. The lowermost segment (SRA; San Angel Limestone [0–97 m]) is exposed in the canyon; the middle segment (SRB; San Angel Limestone and La Penã Formation [97– 209 m]) is located along the highway, and the upper segment (SRC; Tamaulipas Limestone [209–330 m]) is in the canyon. Correlation between the three segments was established using detailed structural measurements. The SRA and SRC segments are continuously exposed and appear to be structurally uncomplicated; segment SRB, however, has a 7 m highly weathered and unexposed gap, between 127 and 134 m, that may correspond to a minor fault or a shale unit that was less resistant to weathering (Fig. 3). We carried out a detailed sampling of all sections. Samples were collected in limestone, marlstone and shale intervals. The SRB section and La Boca Canyon were sampled in more detail, with sample spacing at 0.5–1 m. The SRC section, Canyon Los Chorros and Cienega del Toro were collected in less detail, with sample spacing at ;1.5 m. BIOSTRATIGRAPHY Nannofossil biostratigraphy was carried out on all sections. Nannofossils were observed in smear slides prepared using the technique of Monechi and Thierstein (1985) devised for hard lithologies. All slides were observed in a transmitted light microscope at a magnification of 10003. Detailed biostratigraphic investigations were carried out on samples with the highest nannofossil content and near the ends of species ranges. Relative abundances were assigned to each occurrence based on the number of specimens observed per field of view according to the following scheme: an occurrence was described as abundant if more than 10 specimens were observed in each field; common if one to ten specimens were observed in each field, few if one to nine specimens were observed in ten fields of view and rare if less than one specimen was observed in ten fields of view. Total nannofossil abundance in each sample was described using the same scheme. The nannofossil taxonomy applied in this investigation is standard and discussed and illustrated in Trejo (1975), Deres and Arche´rite´guy (1980), Perch-Nielsen (1985) and Bralower et al. (1994). All taxa identified are listed and problematic taxa are discussed in the Appendix. We use the Lower Cretaceous nannofossil zonation of Thierstein (1973) with the abbreviated nomenclature of Roth (1978). Planktonic foraminiferal biostratigraphy was carried out on the Santa Rosa Canyon section. All observations were made on thin sections prepared from samples every 3 to 5 m. All taxa identified are routine. The zonation of Premoli Silva and Sliter (1995) was used to subdivide the section. The Aptian part of this zonation is based largely on the scheme developed in Mexico by Longoria (1974).

Total Organic Carbon and Carbonate To measure total organic carbon (TOC), samples were decalcified (acid leached in 1N HCl at 258C), concentrated on glass-fiber filters, washed with DDIW, and dried at 508C. TOC was determined by combustion in a LECO C/S 300 analyzer at Indiana University (IU). Precision of the TOC analyses for replicate samples and standards was better than 60.05%. Carbonate analyses were conducted on a Coulometrics coulometer at the University of North Carolina Chapel Hill with a precision of 60.05%.

Carbon Isotopes Carbon isotope measurements were made on the organic carbon fractions of samples. TOC values were used to estimate the amount of powdered sample needed to generate 50 mmol of organically-derived CO2. Samples were placed in 50mL centrifuge tubes with a small amount of DDIW and 6N HCl was added. Samples were centrifuged, and the supernatant removed. This process was repeated until no carbonate reaction was visible. Residues were rinsed in DDIW via centrifuge and dried at 608C. Residue pellets were placed in 9mm annealed quartz tubes, and cupric oxide and copper were added. Samples were evacuated, sealed, and combusted at 8508C. Carbon dioxide was cryogenically distilled. Carbon isotope measurements were conducted at IU on a Finnigan MAT Delta E and at Lehigh University on a Finnigan MAT 252. The C-isotope ratios are expressed on a per mil (‰) basis relative to the Pee Dee Belemnite standard (PDB): d13 C 5 (R sample /R PDB 2 1) 3 1000

where R 5

13

C/ 12 C

Precision of isotopic analyses for replicate samples and lab standards was better than 60.25 per mil.

Rock-Eval Thermal evolution (pyrolysis) was carried out using Rock-Eval analyses performed on eighteen powdered higher-TOC samples from the SRB section by Humble Geochemical Services (Humble, TX). The instrument was calibrated using an internal rock standard. S1 represents the first hydrocarbon peak detected after heating the powdered samples at 3008C for 3 minutes and is proportional to the extractable bitumen in the sample. S2 is the second hydrocarbon peak, produced by the cracking of solid organic matter when the rock is heated from 3008C to 6008C at a rate of 258C/minute. S3 is the carbon dioxide peak measured during the heating interval from 3008C to 3908C and represents the measure of the CO2 produced by pyrolysis of the organic matter in the rock. S1, S2 and S3 are reported in millgrams of hydrocarbon or CO2 per gram of dry rock (Table 1).

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FIGURE 3.

Calcareous nannofossil biostratigraphy of the Santa Rosa Canyon section. A-abundant; C-common; F-few; R-rare.


FIGURE 4.

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Calcareous nannofossil biostratigraphy of the La Boca Canyon section. A-abundant; C-common; F-few; R-rare.

RESULTS BIOSTRATIGRAPHY

Nannofossils Santa Rosa Canyon. Nannofossils are virtually absent in the lower and middle part of the San Angel Limestone. Abundance begins to increase in the uppermost part of this unit, but occurrences remain rare (Fig. 3). Nannofossil abundance increases significantly in the La Penã Formation and remains high into the middle part of the Tamaulipas Limestone allowing precise detection of zonal events including the first occurrence (FO) of Rucinolithus irregularis (base of Zone NC6) between SRB 15 and SRB 16 at 106.7m, the FO of Eprolithus floralis (base of Zone NC7) between SRB 46 and SRB 47 at 142.2m, and the FO of Prediscosphaera columnata (base of Zone NC8) between SRC 4 and SRC 6 at 215.8 m (Table 2). The unexposed interval in the lower part of the section (127 to 134 m) lies in the middle of

Zone NC6. Several useful non-zonal datum levels can also be determined including the last occurrence (LO) of Nannoconus steinmannii between SRB 34 and SRB 34.5 at 124.9 m, the FO of N. truittii between SRB 48 and SRB 49 at 143.1 m, and the LO of Micrantholithus hoschulzii between SRB 85 and SRB 86.5 at 175.05 m. With the exception of the latter datum, which defines the base of Subzone NC7C, none of the other subzonal units of Bralower et al. (1993) can be identified. The nannoconid crisis interval (Erba, 1994) lies between 124.9 m (LO of N. steinmannii) and 142.2 m (abrupt increase in abundance of nannoconids). Preservation deteriorates markedly in the middle part of the Tamaulipas Limestone and nannofossils remain very rare in the upper part of this unit and in the Cuesta del Cura Formation. Only one event can be detected with any certainty, the FO of Eiffellithus turriseiffelii (base of Zone NC10) between SRC 61 and SRC 64 at 301.05 m. The base of Zone NC9 cannot be determined due to

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FIGURE 5.

Calcareous nannofossil biostratigraphy of the Canyon Los Chorros section. C-common; F-few; R-rare.

the absence of the nominate taxon, Axopodorhabdus albianus. The Barremian/Aptian boundary is placed at 100 m at the base of magnetic Polarity Zone MO (Clement et al., in prep.). If the identification of MO is correct, then the true FO of Rucinolithus irregularis should lie somewhat lower in the section than the level determined here (e.g., Coccioni et al., 1992). The Aptian/Albian boundary is placed just above the FO of Prediscosphaera columnata at 220 m. La Boca Canyon. Almost all samples collected from the San Angel Limestone are barren of nannofossils. Nannofossils are rare and preservation is poor in most samples from the La Penã Formation (Fig. 4). The lowermost part of this unit from 208.8 m to 244.15 m correlates with the lower Aptian Chiastozygus litterarius (NC6) Zone based on the presence of Rucinolithus irregularis and the absence of Eprolithus floralis. The assemblage in the sample at 205.8 m is too poorly preserved for the absence of R. irregularis to be interpreted. The absence of Nannoconus steinmannii in this interval indicates correlation to the nannoconid crisis interval in the upper part of Zone NC6 (Erba, 1994). The top of the nannoconid crisis lies at 247.95 m as suggested by a marked increase in the abundance of nannoconids between samples BCA 17 and BCA 18. The FO of Eprolithus floralis lies between samples BCA 15 and BCA 16.1 at 244.15 m, and the E. floralis Zone (NC7) extends to the top of the section (Fig. 4). Other events which can be determined in the La Penã Formation are the LO of Micrantholithus hoschulzii (256.8 m), and the FOs of

Nannoconus truittii (247.95 m) and Prediscosphaera spinosa (287.4 m). Nannoconids are rarer than in the other sections. The Barremian/Aptian boundary is tentatively placed at 200 m below the known occurrence of R. irregularis. Canyon Los Chorros. Nannofossil abundance in samples from Canyon Los Chorros ranges from rare to abundant and preservation from poor to moderate (Fig. 5). A marked increase in abundance and improvement in preservation occurs between the San Angel Limestone and La Penã Formation. The paucity of nannoconids suggests that the base of the section lies within the nannoconid crisis interval. The top of the nannoconid crisis corresponds to the increase in nannoconid abundance between samples LC 4.1 and LC 5.1 (12.5 m). The FO of Rucinolithus irregularis (base of Zone NC6) corresponds to the improvement in preservation at the base of the La Penã Formation (between samples LC 4 and LC 4.1 at 9.15 m). This datum lies above the base of the nannoconid crisis (opposite of other section [Table 2]) and might actually lie somewhat above its true level. The FO of Eprolithus floralis (base of Zone NC7) is placed at 12.5 m (between samples LC 4.1 and LC 5.1). Hence the La Penã Formation correlates with lower Aptian Zone NC6 and upper Aptian Zone NC7. Other events that can be determined include the FO of Nannoconus truittii between samples LC 4.1 and LC 5.1 (12.5 m) and the LO of Micrantholithus hoschulzii (between samples LC 7.1 and LC 7.2; 27.95 m). Cienega del Toro. Nannofossils are exceptionally rare and poorly preserved in most samples from the Cienega


FIGURE 6. Calcareous nannofossil biostratigraphy of the Cienega del Toro section. C-common; F-few; R-rare.

del Toro Section (Fig. 6). The FO of Rucinolithus irregularis (base of Zone NC6) is tentatively placed between CTB 6 and CTB 6.1 (3.72 m). Nannoconus steinmannii has been found in most samples up to CTB 50.21. The absence of this species in the uppermost sample (CTB 50.22) may result from the poor preservation of this sample. The overlap of R. irregularis and N. steinmannii for much or all of this section suggests that it is restricted to the lower Aptian and represents a high rate of sedimentation. Although nannoconids are rare throughout the section, the occurrence of N. steinmannii in the uppermost samples suggests that the section lies below the nannoconid crisis interval (Erba, 1994). The Barremian/ Aptian boundary is tentatively placed at 2 m just below the FO of R. irregularis. Planktic Foraminifers Planktic foraminifers were observed in samples from the Santa Rosa Canyon Section. The abundance of specimens varies significantly; the preservation is predominantly poor. The lowermost part of the SRB section belongs to the Glo-

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bigerinelloides blowi Zone based on presence of the nominate taxon (Fig. 7). The base of the overlying Leupoldina cabri Zone is tentatively placed between samples SRB 43 (139.2 m) and SRB 46 (142.0 m) based on the first occurrence of L. reicheli (Longoria, 1974; Sliter, 1992). However, specimens similar to L. cabri (L. cf. L. cabri), were not observed below sample SRB 49 (143.3 m) and the lowermost L. cabri sensu stricto was observed in sample SRB 54 (146.0 m). The top of the L. cabri Zone, and the base of the overlying Globigerinelloides ferreolensis Zone cannot be defined precisely, and lie between samples SRB 64.5 (154.3 m) (the uppermost specimen of L. cabri) and SRB 84 (172.7 m) (the first occurrence of large, robust specimens of G. ferreolensis). The lowermost occurrence of small, fragile specimens of G. ferreolensis is in sample SRB 70 (159.7 m); the lowermost robust specimens of this species are observed in sample SRB 75 (163.5 m). The overlying Globigerinelloides algerianus Zone ranges from sample SRB 91 (181.1 m) to SRB 107 (197.6 m) based on the occurrence of the nominate taxon. This interval appears to lie in the lower part of the G. algerianus Zone based on the absence of Hedbergella trocoidea. Sample SRB 114 (207.4 m) contains favusellids and calcispheres including Colomiella mexicana and C. recta suggesting a shallow water depositional environment and correlation to the latest Aptian to earliest Albian (Longoria, 1998). Planktic foraminiferal biostratigraphy and assemblages indicate a sequence boundary between samples SRB 107 (197.6 m) and SRB 114 (207.4 m) separating deeper-water and shallow-water facies. The occurrence of favusellids and C. mexicana continues up to sample SRC 15 (230.8 m). However, a prominent dissolution facies is observed between samples SRC 6 (217.3 m) and SRC 10 (223.3 m) that corresponds to OAE1b. Samples SRC 21 (239.8 m) to SRC 56 (292.3 m) appear to correlate to the lower Albian Ticinella primula Zone based on the presence of Hedbergella cf. H. rischi (in sample SRC 21 [239.8 m]) and the nominate taxon (without younger planktic foraminiferal markers) in samples SRC 31 (254.8 m) to SRC 56 (292.3 m). The presence of Clavihedbergella cf. C. simplex in sample SRC 61 (299.8 m) suggests a correlation to the top of the T. primula Zone or the lower part of the Biticinella breggiensis Zone (e.g., Leckie, 1984; Tornaghi and others, 1989; Premoli Silva and Sliter, 1995). The uppermost sample observed, sample SRC 81 (329.8 m), also appears to lie in the lower part of the upper Albian B. breggiensis Zone based on the occurrence of the nominate taxon, T. primula, T. roberti, and the T. roberti group (Leckie, 1984; Tornaghi and others, 1989; Premoli Silva and Sliter, 1995). Geochemistry Organic carbon and carbonate content. Most of the samples analyzed contain low (,0.1%) TOC contents. Three marked peaks (up to 2.7%) in TOC were found in the Santa Rosa Canyon section (Fig. 7), in the lower Aptian, the upper Aptian, and the lowermost Albian. Less marked increases were found in the other sections (Figs. 8–10). Carbonate measurements at Santa Rosa Canyon indicate major facies changes between the San Angel Limestone and the La Penã Formation (in the covered interval between 127

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TABLE 1. Sample

SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB SRB

26.5 34.5 37.5 39.5 40 43.5 46 46.5 56.5 58.5 68.5 82.5 98 99 100 101 106 107

Rock-Eval data from TOC-rich intervals from the Santa Rosa Canyon.

S1

S2

S3

TMAX

TOC (%)

HI

OI

S2/S3

0.01 0.02 0.02 0.01 0.01 0.01 0.02 0.01 0.02 0.01 0.02 0.02 0.00 0.01 0.02 0.03 0.00 0.02

0.05 0.18 0.09 0.04 0.04 0.06 0.13 0.10 0.10 0.00 0.10 0.06 0.02 0.09 0.17 0.13 0.07 0.09

0.42 0.76 0.24 0.65 0.18 0.17 0.23 0.27 0.31 0.38 0.21 0.19 0.32 0.32 0.24 0.17 0.79 0.59

*406 *407 *493 *462 *460 *538 *580 *573 *478 — *388 *328 *428 *427 *574 *524 *379 *380

0.124 0.155 1.183 2.512 1.573 1.394 1.894 1.854 1.316 0.255 0.635 0.287 1.798 1.441 2.664 1.341 1.522 0.497

40.32 116.13 7.61 1.59 2.54 4.30 6.86 5.39 7.60 0.00 15.75 20.91 1.11 6.25 6.38 9.69 4.60 18.11

338.71 490.32 20.29 25.88 11.44 12.20 12.14 14.56 23.56 149.02 33.07 66.20 17.80 22.21 9.01 12.68 51.91 118.71

0.12 0.24 0.38 0.06 0.22 0.35 0.57 0.37 0.32 0.00 0.48 0.32 0.06 0.28 0.71 0.76 0.09 0.15

S1 5 mg. hydrocarbons thermally distilled from 1 gr. of rock. S2 5 mg. of hydrocarbons generated by pyrolytic degradation of the kerogen in one gram of rock. S3 5 mg. of CO2 generated from a gram of rock during temperature programming up to 3908C. TMAX 5 temperature at which maximum S2 hydrocarbons are generated. HI 5 Hydrogen Index (quantity of pyrolyzed organic compounds in S2 relative to TOC). OI 5 Oxygen Index (quantity of CO2 in S3 relative to TOC).

and 134 m). In addition, the three marked peaks in TOC are matched by decreases in CaCO3 content (Fig. 7). Other individual samples with low CaCO3 content represent thin marl and shale seams sampled for micropaleontology. Stable isotopes. C-isotopes measured on the organic fraction (d13Corg) show major stratigraphic fluctuations (Figs. 7– 10). The most detailed, extended record is from the Santa Rosa Canyon section (Fig. 7). This section shows four longterm cyclic fluctuations in d13Corg values. These fluctuations consist of a marked, short-term 0.5–3 per mil decrease in d13Corg values followed by a longer-term 0.5–3 per mil increase. The four decreases are: (1) in the lowermost Aptian nannofossil Zone NC6, top of the Globigerinelloides blowi planktic foraminiferal Zone (139.2 m); (2) in the upper Aptian in the middle of nannofossil Zone NC7, lower part of the Globigerinelloides algerianus planktic foraminiferal Zone (;190 m); (3) in the lowermost Albian at the base of combined nannofossil Zones NC8-NC9, middle part of the combined Hedbergella planispira–Ticinella bejaouaensis Zones (;219 m); and (4) in the upper Albian in the uppermost part of combined nannofossil Zones NC8–NC9 and the top of the Ticinella primula planktic foraminiferal Zone (292–299 m). TABLE 2. Nannofossil event

FO Eiffellithus turriseiffelii FO Prediscosphaera columnata LO Micrantholithus hoschulzii FO Nannoconus truittii Top nannoconid crisis FO Eprolithus floralis LO Nannoconus steinmannii (base nannoconid crisis) FO Rucinolithus irregularis

Carbon isotope stratigraphies of the La Boca Canyon and Canyon Los Chorros sections are limited to the upper Barremian-lower Aptian interval, contain less detail, but show the same basic features (Figs. 8,9) as the contemporaneous part of the Santa Rosa Canyon section. The Cienega del Toro section shows highly fluctuating, but gradually decreasing d13Corg values for most of the section followed by increasing d13Corg values (Fig. 10). This pattern appears to correlate to the lowermost cycle at Santa Rosa Canyon. Rock-Eval. In general, the low organic carbon content of the SRB samples resulted in S1 and S2 values that were too low for confident interpretation (Table 1). As a result, the Tmax values for the samples are not reliable. Likewise, the hydrogen index (HI) and oxygen index (OI) values for these samples plot close to the origin, revealing little about the source or thermal history (Fig. 11). DISCUSSION CHRONOLOGY

APTIAN-ALBIAN C-ISOTOPE EXCURSION

OF THE

The classic paper by Scholle and Arthur (1980), based in part on investigation of the Peregrina Canyon and Rancho

Meter levels of nannofossil datums in sections investigated. Base of zone

Santa Rosa Canyon

La Boca Canyon

Canyon Los Chorros

NC7

301.05 215.80 175.05 143.10 142.20 142.20

256.80 247.95 247.95 244.15

27.95 12.50 12.50 12.50

NC6

124.90 106.70

base base

base 9.15

NC10 NC8

Cienega del Toro

top 3.72


427

FIGURE 7. Carbonate, total organic carbon (TOC), and d13Corg stratigraphy of the Santa Rosa Canyon section. Gaps in the record indicate sampling gaps. Nannofossil zones are from Roth (1978). Planktic foraminiferal zones are from Premoli Silva and Sliter (1995). Shaded area in nannofossil stratigraphy represents the nannoconid crisis interval.

Jacalitos sections in Mexico, demonstrated the applicability of C-isotopes in regional and global correlation. The C-isotope measurements of Scholle and Arthur (1980) were made on the carbonate (d13Ccarb) fraction (Fig. 12). Chronostratigraphic control of the two sections was based largely on unpublished proprietary microfossil and ammonite biostratigraphy. However, the Peregrina Canyon section is almost entirely barren of nannofossils (T. J. Bralower, unpublished data). Subsequent C-isotope stratigraphic studies of the Barremian-Albian interval (e.g., Weissert and others, 1985; 1998; Pratt and King, 1986) have reproduced the basic shape of the Scholle and Arthur (1980) curve, but have revealed numerous differences in its chronostratigraphic correlation. The results of the present investigation allow detailed correlation between nannofossil and foraminiferal biostratigraphy and C-isotope stratigraphy for the late Barremian to late Albian interval. Carbon isotopes, measured on the organic carbon fraction, show identical stratigraphic changes to detailed d13Ccarb and d13Corg curves from European, Middle Eastern and Pacific shallow-water and deep-sea sections (e.g., Weissert and others, 1985; Jenkyns, 1995; Erbacher and others, 1996; Vahrenkamp, 1996; Ferreri and others, 1997; Menegatti and others, 1998; Gro¨cke and others, 1999)

(Fig. 13). Most detailed curves from Europe and the Pacific terminate in the late Aptian or early Albian. The only curve that extends through the late Albian is that of Erbacher and others (1996). The present investigation adds new detail to the Erbacher and others (1996) curve and to the C-isotope stratigraphy of the middle and late Albian interval. High-resolution Aptian d13Ccarb and d13Corg stratigraphies of sections from Alpine Tethys show several distinct segments (Menegatti and others 1998). These include: a significant decrease in d13C values (segment C3), followed by a marked increase (C4), an interval of level d13C values (C5), further increase (C6), then somewhat variable but overall constant values (C7), followed by decreasing values (C8) (Fig. 13). In the Tethyan sections, segment C3 to lower segment C7 lie in the Chiastozygus litterarius nannofossil Zone (NC6) and the Globigerinelloides blowi planktic foraminiferal Zone (Menegatti and others, 1998), and the remainder of the segments lie in the Rhagodiscus angustus nannofossil Zone (NC7). The upper part of segment C7 lies in the upper part of the Leupoldina cabri Zone and segment C8 begins in the L. cabri Zone and continues into the Globigerinelloides algerianus Zone (Fig. 13). Carbon-isotope stratigraphy of the four Mexican sec-

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FIGURE 8. Total organic carbon (TOC), and d13Corg stratigraphy of the La Boca Canyon section. Nannofossil zones are from Roth (1978). Shaded area in nannofossil stratigraphy represents the nannoconid crisis interval.

tions show many of the same features as the Menegatti and others (1998) curve, thus we have adopted the terminology of these authors and extend it to the late Albian using the record at Santa Rosa Canyon (Fig. 7). This record shows a sharp increase in d13Corg values (segment C9) in the G. algerianus planktic foraminiferal Zone and Rhagodiscus angustus nannofossil Zone, followed by level d13Corg values (segment C10) for the upper parts of the same zones extending into the combined Hedbergella planispira–Ticinella bejaouaensis planktic foraminiferal Zones, sharply decreasing (C11), then sharply increasing (C12) d13Corg values in the middle and upper part of the combined H. planispira–T. bejaouaensis planktic foraminiferal Zones and the lower part of the combined Prediscosphaera columnata (NC8)–Axopodorhabdus albianus (NC9) nannofossil Zones, then fairly constant d13Corg values that extend to the top of the Ticinella primula foraminiferal Zone and the combined P. columnata–A. albianus nannofossil Zones (segment C13). Finally, a sharp decrease in d13Corg values

(segment C14) occurs at the top of the T. primula Zone and the combined P. columnata–A. albianus nannofossil Zones, and a slight increase in d13Corg values (segment C15) occurs in the lower parts of the Biticinella breggiensis foraminiferal and Eiffellithus turriseiffelii nannofossil Zones. The Santa Rosa Canyon d13Corg curve shows several of the same features as the d13Ccarb curve from Peregrina Canyon (Scholle and Arthur, 1980), but with a different chronostratigraphic correlation (Fig. 12). Since the early Aptian negative-positive d13C excursions (segments C3 to C7) are distinct, we use these features to anchor the Peregrina Canyon d13Ccarb curve and tentatively identify each of the Csegments identified at Santa Rosa Canyon (Fig. 7). Using the correlation with Santa Rosa Canyon, we have moved the Barremian/Aptian and the Aptian/Albian boundaries upwards at Peregrina Canyon (Fig. 12). The similarity of C-isotope curves and a range of biochronologic data (e.g., Premoli Silva and others, 1989; Erba,


429

FIGURE 9. Total organic carbon (TOC), and d13Corg stratigraphy of the Canyon Los Chorros section. Nannofossil zones are from Roth (1978). Shaded area in nannofossil stratigraphy represents the nannoconid crisis interval.

1991; Herbert, 1992; Erba, 1996) suggest that sedimentation rates in Italian Barremian-Aptian sections with d13C stratigraphies (i.e., Cismon, Piobbico) were fairly constant. Carbon isotope data and general paleogeographic factors indicate that the Mexican sections were characterized by rather dramatic changes in sedimentation rates and that differences between sections were marked. For example, the Cienega del Toro section had much higher overall sedimentation rates during the earliest Aptian than the other sections. However, particular details of the Barremian-Albian C-isotope records of the hemipelagic Mexican sections and pelagic sections from southern Europe and the deep sea are remarkably similar. Sufficient detail exists in the C-isotope record of this time interval (Fig. 13) that C-isotope chemostratigraphy can provide precise age control in shallowwater limestone sections with poor biostratigraphic control (e.g., Fo¨llmi and others, 1994; Jenkyns, 1995; Vahrenkamp, 1996; Lehmann and others, in press).

ORIGIN

OF

APTIAN-ALBIAN C-ISOTOPE VARIATIONS

Effect of Corg Source and Diagenesis on d13Corg Values Variations in d13Corg values can result from a variety of different factors. These include changes in the source of Corg (e.g., Arthur and others 1985b), and selective diagenetic alteration of particular samples. Before the paleoenvironmental significance of d13Corg stratigraphies can be established, the significance of source and diagenetic changes must be established. The Rock-Eval data (Fig. 11) reveal little about the source of Corg. However, major changes in source seem unlikely given the hemipelagic and pelagic depositional environment of the sequences investigated. Although the Santa Rosa Canyon section was sampled in more detail than the other three sequences, and hence the C-isotope stratigraphy of this section is more detailed, the sections can be correlated with one another and with C-isotope stratigraphies of other sections (Fig. 13). Thus, although absolute d13Corg val-

430

FIGURE 10.


Total organic carbon (TOC), and d13Corg stratigraphy of the Cienega del Toro section. Nannofossil zones are from Roth (1978).

ues may have been altered during burial, diagenesis does not appear to be responsible for the major trends observed.

Carbon Isotope Shifts and Aptian-Albian Oceanic Anoxic Events The Santa Rosa Canyon d13Corg curve (Fig. 7) shows an interesting relationship with oceanic anoxic events, early Aptian event OAE1a, earliest Albian event OAE1b (Arthur and others, 1990; Bralower and others, 1993), and a new event in the late Aptian. All three events correspond to short-term, 0.5–3 per mil decreases in d13Corg values followed by increases of similar magnitude. The increases in d13Corg values and the long-term Aptian-Albian excursion likely reflect the burial of isotopically light Corg from the marine reservoir (e.g., Scholle and Arthur, 1980; Arthur and others, 1985a; Weissert and others, 1985; Weissert and others, 1998). The negative excursions have previously been observed in high-resolution records (e.g., Jenkyns, 1995; Menegatti and others, 1998; Gro¨cke and others, 1999; Erba and others, 1999) and their origin is not fully understood.

A change in carbon cycling (e.g., Menegatti and others, 1998), an increase in hydrothermal activity (e.g., Arthur and others, 1991; Bralower and others, 1994; Leckie and others, 1998), or thermal dissociation of methane hydrates (Jahren and Arens, 1998) are possible explanations. Oceanic anoxic events OAE1a (segments C3 to C6) and OAE1b (segments C11 to C12) appear to be global in extent (e.g., Bralower and others, 1993). This is suggested by: (1) the occurrence of Corg-rich horizons in the Santa Rosa Canyon corresponding to levels identified as OAE1a and OAE1b in widespread European land sections and deep-sea sections, and (2) the similar negative-positive response of C-isotope curves from Santa Rosa Canyon and other sections. The lack of Corg-rich horizons or significant C-isotope anomalies in the Axopodorhabus albianus nannofossil Zone in Santa Rosa Canyon suggests that OAE1c (Bralower and others, 1993; Erbacher and others, 1996) was more regional in extent. The late Aptian event (segments C8 and C9) lies in the long duration Rhagodiscus angustus nannofossil Zone


431

FIGURE 11. Plot of oxygen versus hydrogen indices for select samples across organic-rich intervals at Santa Rosa Canyon.

(NC7) and the upper part of the Globigerinelloides algerianus planktic foraminiferal Zone (Fig. 7). Segments C8 and C9 can been recognized in other C-isotope records (e.g., Weissert and Lini, 1991; Jenkyns, 1995; Erbacher and others, 1996; Vahrenkamp, 1996; Ferreri and others, 1997; Weissert and others, 1998; Menegatti and others, 1998; Gro¨cke and others, 1999) (Fig. 12), however neither segment in the other sections corresponds to Corg-rich sediments. The Corg-rich interval in Santa Rosa Canyon lies stratigraphically above a Corg-rich level in the Calera Limestone in the Globigerinelloides ferreolensis Zone (Sliter, 1989); it lies below the Corg-rich Jacob level and Level 113 which are observed in the Ticinella bejaouaensis planktic foraminiferal Zone in sequences from the Fosse Vocontien, southern France, and the Umbrian Apennines of Italy, respectively (e.g., Tornaghi and others, 1989; Weissert and Bre´he´ret, 1991). The late Aptian decrease in d13C values has previously been assigned to an interval of cooling, lower sea level (e.g., Haq and others, 1987) and reduced Corg burial (Weissert and Lini, 1991). However, this interval in Mexico clearly corresponds to increased Corg contents (Fig. 7) and rising relative sea-level (Lehmann and others, in press). The absence of Corg-rich horizons in pelagic sequences from southern Europe indicates that the late Aptian event is not global in extent, however, the negative-positive C-isotope response suggests that Corg burial may have been widespread along continental margins. Alternatively, the increase in Corg content in the Santa Rosa Canyon section may be a local phenomenon caused by a decrease in carbonate accumulation. Clearly, the late Aptian event warrants further investigation.

Carbon Isotope Variations: Understanding Global Controls Understanding the primary oceanographic, eustatic and basinal processes responsible for detailed Aptian-Albian C-

FIGURE 12. Barremian-Albian C-isotope stratigraphy of the Peregrina Canyon section after Scholle and Arthur (1980). C-segments C1 to C8 after Menegatti et al. (1998), C9 to C15 as defined in Santa Rosa Canyon (Fig. 7). Stage boundaries determined by Scholle and Arthur (1980) and revised in this investigation are shown at right (see text for details).

isotope fluctuations (Fig. 13) is difficult. Originally such variations were thought to result primarily from increased burial and subsequent exhumation and oxidation of Corg during OAEs (e.g., Scholle and Arthur, 1980). More recently, however, additional factors have been proposed. These include input of isotopically light (26 to 7 per mil) volcanic CO2 (e.g., Arthur and others, 1991; Caldeira and Rampino, 1991; Bralower and others, 1994), increased recycling rates of 12C-rich intermediate water (e.g., Menegatti and others, 1998), intensified flux of 12C-rich riverine DIC (e.g., Weissert, 1989), and thermal dissociation of methane hydrate (Jahren and Arens, 1998). Since several of these mechanisms are potentially at work, the causes of Aptian-Albian d13C fluctuations are clearly complex (e.g., Menegatti and others, 1998). At least part of the negative C-isotope excursion that precedes OAE1a can be explained by volcanism; the emplace-

432


FIGURE 13. Correlation of carbon isotope stratigraphies from the Santa Rosa Canyon section (d13Corg), Peregrina Canyon section, Mexico (d13Ccarb after Scholle and Arthur, 1980), Cismon section, S. Alps, Italy (d13Ccarb, after Menegatti et al., 1998), DSDP Site 463, Mid-Pacific Mountains (d13Corg from D. Allard, unpubl. data), and ODP Site 866, Resolution Guyot (d13Ccarb after Jenkyns, 1995). Depth scales show tick marks every 20 m for Cismon, Site 463, and Santa Rosa Canyon, and 100 m for Site 866 and Peregrina Canyon.

ment of the massive Ontong Java LIP is synchronous with the isotopic shift and sea-floor spreading rates also increased at this time (Arthur and others, 1991; Larson, 1991; Erba, 1994; Bralower and others, 1994). Although the correlation is less precise, OAE1b possibly correlates with the emplacement of the Kerguelen LIP in the early Albian (e.g., Leckie and others, 1998). No known LIP episode correlates with the late Aptian C-isotope event. Explaining the negative C-isotope excursion by thermal dissociation of methane hydrate (e.g., Jahren and Arens, 1998) is attractive from a volumetric point of view. The Cisotopic composition of methane is so negative (;260–65 per mil) that a small proportion of the global methane reservoir can cause a sizable negative global excursion (e.g., Dickens and others, 1995). Given the warm temperatures of mid-Cretaceous intermediate and deep waters (e.g., Douglas and Savin, 1975; Huber and others, 1995; Fassell and Bra-

lower, 1999), however, it is unlikely that a significant methane hydrate reservoir existed at this time. Variations in carbon isotopes measured on the organic fraction of Aptian-Albian sediments are larger than those measured on the carbonate fraction; this is thought to reflect increased atmospheric CO2 contents (e.g., Hayes and others, 1989; Popp and others, 1989; Menegatti and others, 1998). For example, at Cismon (Southern Alps, Italy) d13Corg variations are ;7 per mil, d13Ccarb variations are ;3 per mil; at Rotter Satel (Swiss Alps) d13Corg variations are ;5 per mil, d13Ccarb variations are ;3 per mil (Menegatti and others, 1998); at ODP Site 866 on Resolution Guyot, d13Corg variations are .6 per mil, d13Ccarb variations are ;4.5 per mil (Jenkyns, 1995; Baudin and Sachsenhofer, 1996). In Mexico, the size of carbon isotopic fluctuations is smaller than in these other locations: d13Corg variations are ;3 per mil and d13Ccarb variations are ;2 per mil (Scholle and Arthur,

433


1980; Fig. 12). We speculate that the amplitude of the Mexican d13Corg variations has been decreased by diagenetic alteration accompanying burial and subsequent tectonic uplift of the Sierra Madre. Menegatti and others (1998) showed that the decrease in d13C values and the increase in D13C (d13Corg-d13Ccarb) preceded the onset of elevated Corg contents in OAE1a. Such a relationship is not observed for any of the OAEs in the Santa Rosa Canyon Section (Fig. 7). However, the base of OAE1a is contained within the covered interval and thus the relative timing of the decrease in d13Corg values and the increase in Corg contents cannot be observed. In addition, level d13C values that characterize segment C5 in the records of Menegatti and others (1998) cannot be observed in the Mexican sections. One possibility is that this segment also lies within the covered interval and that segments C3 and C4 have been misidentified. CLUES

TO THE

ORIGIN

OF

OCEANIC ANOXIC EVENTS

Potential mechanisms for increased burial of Corg in the Aptian-Albian OAE were discussed by Weissert (1989). He proposed that increased flux of hydrothermal CO2 into the atmosphere caused warm and humid conditions which led in turn to more intensive continental weathering and increased runoff. Heightened continental runoff increased phosphorus and dissolved inorganic carbon flux to the oceans, which increased primary productivity. Warm and fresh surface waters led to increased water-column stability and deep-water stagnation that caused increased burial rates of Corg in sediments. Increased sea floor spreading rates and mid-plate volcanism (e.g., Larson, 1991; Tarduno and others, 1991) also led to eustatic rise in sea level which would have increased the area of upwelling along continental shelves. Although it is difficult to understand the relative role of all of the different factors cited by Weissert (1989), subsequent authors (e.g., Arthur and others, 1990; Weissert and Lini, 1991; Bralower and others, 1994; Menegatti and others, 1998) have proposed similar scenarios for the Aptian-Albian OAE. One major problem with many interpretations of pelagic sections is that the correlation of environmental changes with relative changes in sea level is indirect and thus subject to considerable error. The Mexican sections investigated can be directly correlated with shallow-water sequences in which relative sea level changes have been interpreted. An early Aptian sea level rise has been proposed to have drowned the Cupido Platform to the northwest of the study area (e.g., Goldhammer and others, 1991). However, peak flooding, associated with the deposition of the La Penã Formation and termination of the platform, did not occur until the mid to late Aptian (Lehmann and others, in press). Platforms in Europe and the Middle East (Fo¨llmi and others, 1994; Vahrenkamp, 1996; Ferreri and others, 1997; Weissert and others, 1998) and atolls in the Pacific (e.g., Ro¨hl and Ogg, 1996) show a similar relative sea level chronology. The transition from the San Angel Limestone to the La Penã Formation at Santa Rosa Canyon is associated with an overall decrease in CaCO3 content (Fig. 7), but an increase in the abundance of planktic foraminifers and coccoliths (nannofossils excluding nannoconids). These changes may

reflect the combined demise of the nannoconids, prolific carbonate producers, and a decrease in dilution of carbonate detritus derived from the platform to the west as a result of a relative sea level rise. Although sea level was clearly rising, OAE1a appears to have taken place before peak flooding and before its associated condensed section (Loutit and others, 1989). A similar relationship has been observed by Fo¨llmi and others, (1994), Vahrenkamp (1996) and Weissert and others (1998). In Mexico, the late Aptian anoxic event and OAE1b appear to lie in intervals of rising relative sea level, although subdivision of the transgressive part of the sea level cycle is not possible due to the lack of lithologic variability in the La Penã Formation (Lehmann and others, in press). Other investigations show variable interpretations of relative sea level change in these time intervals: Fo¨llmi and others (1994) and Vahrenkamp (1996) show generally rising sea level; Weissert and others (1998) show generally falling relative sea level. Thus the three anoxic episodes documented in the Mexican sections appear to lie somewhere within the transgressive systems tract suggesting a constant relationship with an external forcing mechanism, possibly volcanism (e.g., Arthur and others, 1985; Bralower and others, 1994). THE APTIAN ‘‘NANNOCONID CRISIS’’

IN

MEXICO

The ‘‘nannoconid crisis’’ observed in lowermost Aptian sediments from Europe, and the Atlantic and Pacific Oceans (Coccioni and others, 1992; Erba, 1994; Bralower and others, 1994) involves a dramatic reduction in the number of narrow-canal nannoconid taxa (including Nannoconus colomii and N. steinmannii), a short interval with few or no nannoconids and high abundances of other taxa such as Assipetra infracretacea and Rucinolithus terebrodentarius (the crisis interval), followed by an increase in the number of nannoconid taxa with wide canals (N. globulus, N. bucheri, N. truittii, and N. kamptneri). In the European and Pacific sections, the ‘‘nannoconid crisis’’ occurs within nannofossil Zone NC6, just above magnetic Polarity Zone M0 and below organic-rich sediments of OAE1a (Erba, 1994). In Deep Sea Drilling Project (DSDP) Site 641, however, the crisis occurs slightly earlier, at the base of Polarity Zone M0 (Bralower and others 1994). The significance of this minor diachroneity is not understood. Although poor preservation in the Mexican samples has reduced the number of nannoconids, it is still possible to detect the extinction of narrow-canal nannoconids and the appearance of the wide-canal taxa. The top of the nannoconid crisis is much easier to detect than the base, corresponding to a sharp increase in the relative abundance of nannoconids in the Santa Rosa Canyon, La Boca Canyon, and Canyon Los Chorros sections. The ‘‘crisis’’ interval lies between 124.9 m and 142.2 m in the Santa Rosa Canyon section, between the base of the sampled section (205.8 m) and 247.95 m in the La Boca Canyon section and between the base of the sampled section and 12.5 m in the Los Chorros Canyon section. The stratigraphic distribution of nannoconid taxa determined here is similar to that described by Trejo (1975) in some of these same sections. We have found rare specimens of Nannoconus wassallii, N. circularis, N. bucheri, and N. kamptneri

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in the crisis interval. The diversity of nannoconids thus appears to increase. While we do observe Rucinolithus terebrodentarius and Assipetra infracretacea in the crisis interval, these species are in much lower abundance than described by Erba (1994) in Tethyan and Pacific sections. The crisis interval is best preserved in the Santa Rosa Canyon section. Here the sequence of biohorizons is as follows: LO of N. steinmanii, FO of N. wassallii, FO of N. truittii, and FO of N. vocontiensis. In all of the sections, the top of the crisis interval is close to the FO of Eprolithus floralis (base of Zone NC7) and to the FO of N. truittii. The timing of the crisis interval relative to nannofossil biohorizons and magnetic Polarity Zone M0 is almost identical to the results of Erba (1994). CONCLUSIONS Carbon isotope stratigraphy from four Mexican sections shows similar stratigraphic changes to curves from Barremian to lower Albian European and Pacific deep-sea sections. Three abrupt peaks in Corg content correlate with OAE1a (early Aptian), OAE1b (earliest Albian) and a new event in the late Aptian Globigerinelloides algerianus foraminiferal Zone. All three events correspond to short-term, 0.5–3 per mil decreases in C-isotope values followed by increases of similar magnitude. The similar shape of the Cisotope curves in Mexico and other areas, and the response of C-isotopes to the OAEs, indicate that the late Aptian episode was extensive and that OAE1a and OAE1b were global. The three anoxic episodes appear to lie within the transgressive systems tract suggesting a constant relationship with an external forcing mechanism such as volcanism. The early Aptian ‘‘nannoconid crisis’’ is observed in all of the sections. ACKNOWLEDGMENTS We thank Tim Herbert, Christoph Lehmann, Isabel Montan˜ez, David Osleger, and Chris Burnett for field assistance, and Mark Leckie and Helmi Weissert for reviewing the initial manuscript. We are extremely grateful to Lisa Pratt (Indiana University) and Gray Bebout (Lehigh University) for allowing us to use their analytical facilities and for assistance with methods. We thank Melissa Amentt for running coulometer analyses, Peter Scholle, Hugh Jenkyns and Helmi Weissert for making available their published isotope data sets, and Christoph Lehmann for sharing unpublished sea level interpretation. All geochemical data are archived at the World Data Center-A for Paleoclimatology, NOAA/ NGDC 325 Broadway, Boulder, CO 80303 (phone 303-4976280; e-mail: [email protected]). This research was supported by NSF-EAR-9305727 to Bralower and CoBabe and NSF-EAR-9313914 to Clement and Longoria. REFERENCES ARTHUR, M. A., and PREMOLI SILVA, I., 1982, Development of widespread organic carbon-rich strata in Mediterranean Tethys, in Schlanger, S. O., and Cita, M. B. (eds.), Nature and Origin of Cretaceous Carbon-Rich Facies, Academic Press, London: p. 7– 54. , DEAN, W. E., and SCHLANGER, S. O., 1985a, Variations in global carbon cycling during the Cretaceous related to climate,

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Received 1 February 1999 Accepted 8 June 1999 APPENDIX: NANNOFOSSIL TAXONOMY

Lithraphidites alatus magnus Covington and Wise, 1987 Hayesites albiensis Manivit, 1971 Parhabdolithus achlyostaurion Hill, 1976 Corollithion achylosum (Stover, 1966) Thierstein, 1971 Rhagodiscus angustus (Stradner, 1963) Bralower, Erba and Mutterlose in Bralower et al., 1994 Rhagodiscus asper (Stradner, 1963) Manivit, 1971 Watznaueria barnesae (Black, 1959) Perch-Nielsen, 1968 Nannoconus bermudezii Bro¨nnimann, 1955 Watznaueria biporta Bukry, 1969 Nannoconus bonetii Trejo, 1959 Watznaueria britannica (Stradner, 1963) Reinhardt, 1964 Nannoconus bucheri Bro¨nnimann, 1955 Lithraphidites carniolensis Deflandre, 1963 Microstaurus chiastius (Worsley, 1971) Bralower et al., 1989 Nannoconus circularis Deres and Acheriteguy, 1980 Markalius circumradiatus (Stover, 1966) Perch-Nielsen, 1968 Prediscosphaera columnata (Stover, 1966) Perch-Nielsen, 1984 Watznaueria communis Reinhardt, 1964 Cretarhabdus conicus Bramlette & Martini, 1964 Biscutum constans (Gorka, 1957) Black ex Black & Barnes, 1959


Grantarhabdus coronadventis (Reinhardt, 1966) Gru¨n in Gru¨n and Allemann, 1975 Tetrapodorhabdus decorus (Deflandre, 1954) Wind and Wise in Wise and Wind, 1977 Zygodiscus diplogrammus (Deflandre & Fert, 1954) Gartner, 1968 Cribrosphaerella ehrenbergii (Arkhangelsky, 1912) Deflandre in Piveteau, 1952 Zygodiscus elegans (Gartner, 1968) Bukry, 1969 Parhabdolithus embergeri (Noe¨l, 1959) Bralower, Monechi & Thierstein, 1989 Zygodiscus erectus (Deflandre, 1954) Bralower, Monechi & Thierstein, 1989 Percivalia fenestrata (Worsley, 1971) Wise, 1983 Eprolithus floralis (Stradner, 1962) Stover, 1966 Tranolithus gabalus Stover, 1966 Nannoconus globulus (Bro¨nnimann, 1955) subsp. globulus Micrantholithus hoschulzii (Reinhardt, 1966) Thierstein, 1971 Assipetra infracretacea (Thierstein, 1973) Roth, 1973 Rucinolithus irregularis Thierstein in Roth and Thierstein, 1972 Nannoconus kamptneri (Bro¨nnimann, 1955) subsp. kamptneri Rotelapillus laffittei (Noe¨l, 1956) Noe¨l, 1973 Diazomatolithus lehmanii Noe¨l, 1965 Chiastozygus litterarius (Gorka, 1957) Manivit, 1971 Cyclagelosphaera margerelii Noe¨l, 1965 Eiffellithus monechii (Hill and Bralower, 1987) Crux, 1991 Flabellites oblongus (Thierstein, 1973) Crux, 1982 Micrantholithus obtusus Stradner, 1963

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Tranolithus orionatus (Reinhardt, 1966) Perch-Nielsen, 1968 Watznaueria ovata Bukry, 1969 Manivitella pemmatoidea (Deflandre ex Manivit, 1965) Thierstein, 1971 Braarudosphaera regularis Black, 1973 Bidiscus rotatorius (Bukry, 1969) Thierstein, 1973 Cretarhabdus schizobrachiatus (Gartner, 1968) Bukry, 1969 Micrantholithus speetonensis Perch-Nielsen, 1979 Prediscosphaera spinosa (Bramlette and Martini, 1964) Gartner, 1968 Rhagodiscus splendens (Deflandre, 1953) Noe¨l, 1969 Nannoconus steinmannii (Kamptner, 1931) subsp. steinmannii. We combine this species with N. colomii as it is not always possible to discern a central cavity. Tegumentum stradneri Thierstein, in Roth & Thierstein, 1972 Vagalapilla stradneri (Rood, Hay & Barnard, 1971) Thierstein, 1973 Tegumentum striatum (Black, 1971) Taylor, 1979 Cretarhabdus surirellus (Deflandre, 1954) Reinhardt, 1970 Rucinolithus terebrodentarius Applegate et al., in Covington & Wise, 1987 Eiffellithus trabeculatus Gorka, 1957 Eiffellithus turriseiffelii (Deflandre in Deflandre and Fert, 1954) Reinhardt, 1965 Nannoconus truittii Bro¨nnimann, 1955 Nannoconus vocontiensis Deres and Ache´rite´guy, 1980 Nannoconus wassallii Bro¨nnimann, 1955 Chiastozygus spp. Nannoconus spp.