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Turonian paleoceanographic event: the Goban Spur stable isotope record
Palaeogeography, Palaeoclimatology, Palaeoecology 201 (2003) 51^66 www.elsevier.com/locate/palaeo
Changes in Northeast Atlantic temperature and carbon £ux during the Cenomanian/Turonian paleoceanographic event: the Goban Spur stable isotope record M. Gustafsson , A. Holbourn, W. Kuhnt Institut fu«r Geowissenschaften, Christian-Albrechts-Universita«t, Olshausenstrasse 40, D-24118 Kiel, Germany Received 20 November 2001; received in revised form 2 June 2003; accepted 17 June 2003
Abstract Stable isotopes of bulk sediment and well preserved tests of planktonic and benthic foraminifera from midlatitude NE Atlantic DSDP Site 551 (Goban Spur) provide the first estimates of carbon isotope gradients within the water column at a lower bathyal site during the Cenomanian/Turonian boundary interval (CTBI). The CTBI carbon isotope excursion is prominent (up to 2x shift in N13 C) in the bulk (coccolith) signal, but less pronounced (approximately 0.5x shift in N13 C) in planktonic and benthic foraminifera. This difference indicates a very steep 13 C gradient in the upper water column and a very efficient biological pump during the CTBI carbon isotope excursion. We suggest significantly increased seasonal primary production in the uppermost water column with an enhanced shallow water chlorophyll maximum as a cause for this steep carbon isotope gradient. Deep-water and surface-water temperature changes during the CTBI are estimated using benthic and planktonic foraminiferal oxygen isotopes. Warm deep-water masses (13^16‡C) and a low temperature gradient within the water column prevailed in the late Cenomanian. Additional warming (approximately 2‡C for both surface and deep water) occurred in the latest Cenomanian prior to CTBI black shale deposition. This pattern of CTBI black shale deposition during a temperature maximum is also evident at two low latitude locations (ODP Site 1050, Blake Nose and Tarfaya, southern Morocco). < 2003 Published by Elsevier B.V. Keywords: Cenomanian^Turonian boundary; stable isotopes; foraminifer; carbon cycle; NE Atlantic; paleotemperature
1. Introduction The mid-Cretaceous represents one of the most remarkable episodes of greenhouse climate in Earth’s history. It was characterized by high atmospheric CO2 concentrations, low latitudinal temperature gradient and unusually high deep-
* Corresponding author. Fax: +49-431-880-4376. E-mail address: [email protected] (M. Gustafsson).
water temperatures in the global ocean (Hay, 1995). The mid-Cretaceous greenhouse was coincident with a worldwide pulse in ocean crustal production (Larson, 1991). The release of mantle CO2 from this very active volcanic episode may have in fact directly caused the warm mid-Cretaceous greenhouse climate (Larson, 1991). Thus, the study of mid-Cretaceous paleoclimate and paleoceanography provides insight into a natural climatic experiment, when a large amount of CO2 was released into the atmosphere. During
0031-0182 / 03 / $ ^ see front matter < 2003 Published by Elsevier B.V. doi:10.1016/S0031-0182(03)00509-1
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Fig. 1. Paleogeographical map of the mid-Cretaceous (92 Ma) North Atlantic showing the position of DSDP Site 551, ODP Site 1050, Drill Hole S75 and the Eastbourne section. Present day shorelines are indicated and gray shaded areas represent continental reconstruction from data ¢les in Hay et al. (1999), downloaded from http://www.odsn.de/odsn/services/paleomap.
the mid-Cretaceous, widespread deposition of organic carbon occurred during several oceanic anoxic events (OAEs). Erbacher et al. (1999) de¢ned ¢ve OAEs, and the Cenomanian/Turonian boundary interval (CTBI) represents one of the more severe and long-lasting of those events. Caron et al. (1999) estimated the duration of the CTBI as 0.4 myr., but both longer and shorter durations have been proposed by other authors. A positive carbon isotope excursion of up to 2x in carbonate carbon and up to 6x in organic carbon is associated with the CTBI (Arthur et al., 1988). Most likely, the carbon isotope excursion resulted from the removal of light 12 C from the global reservoir by the globally enhanced burial of 13 Cdepleted organic carbon into black shales (Arthur et al., 1988). However, it is still a matter of dis-
cussion whether this increased burial of organic carbon was mainly caused by high primary production (e⁄cient biological pump) or by enhanced organic matter preservation at comparatively low primary production levels. We present new oxygen and carbon isotope data across the CTBI from Deep Sea Drilling Program (DSDP) Site 551, which is located on the Goban Spur in the eastern North Atlantic (de Graciansky and Bourbon, 1985). The unusually good preservation of calcareous foraminifera and a paleodepth close to, but above the Calcite Compensation Depth a¡ord a unique opportunity to estimate both the carbon isotope gradient within the water column and the oxygen isotopic paleotemperatures of deep and surface waters. The isotopic results are compared with those pub-
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Fig. 2. Bulk sediment carbon isotope (N13 C) curve and lithological summary for Site 551 (modi¢ed after de Graciansky and Bourbon, 1985).
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Table 1 Oxygen and carbon isotope data for bulk sediment, fraction 6 63 Wm, planktonic foraminifera and benthic foraminifera from Site 551. Sample 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc, 80-551-5cc, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2,
6^9 16^19 27^30 39^42 39^42 48^51 67^70 67^70 78^81 87^90 96^99 96^99 108^111 116^119 127^130 127^130 139^142 139^142 146^150 6^9 18^21 18^21 18^21 27^30 37^40 48^50 57^60 76^79 86^89 96^99 106^109 106^109 116^118 4^7 13^16 17^20 37^40 37^40 58^61 82^85 97^100 117^120 117^120 139^141 17^20 17^20 37^40 57^60 57^60 76^79 76^79 76^79 97^100
MBS
Bulk samples
N13 C
N18 O
Sample
132.56 132.66 132.77 132.89 132.89 132.98 133.07 133.07 133.28 133.37 133.46 133.46 133.58 133.66 133.77 133.77 133.89 133.89 133.96 134.06 134.18 134.18 134.18 134.27 134.37 134.48 134.57 134.76 134.86 134.96 135.06 135.06 135.16 135.54 135.63 138.67 138.87 138.87 139.08 139.32 139.47 139.67 139.67 139.89 140.17 140.17 140.37 140.57 140.57 140.76 140.76 140.76 140.97
Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment
2.12 2.77 2.84 3.18 3.12 2.84 2.64 2.63 2.49 2.73 2.64 2.66 2.64 2.75 2.66 2.66 2.70 2.69 2.69 3.18 2.90 2.89 2.93 2.94 3.05 3.18 3.80 3.92 3.68 3.52 3.01 3.01 3.67 3.43 3.28 2.28 2.43 2.57 2.33 2.58 2.43 2.33 2.43 2.35 2.29 2.35 2.20 2.25 2.28 2.16 2.09 2.22 2.40
31.49 31.88 31.89 31.63 31.52 32.56 32.08 31.93 32.27 32.09 32.36 32.11 32.12 31.90 32.33 32.28 32.04 32.07 32.14 32.63 31.88 32.09 31.90 32.16 31.79 31.89 32.57 32.92 32.59 32.57 32.29 32.20 32.50 32.46 32.22 31.39 31.38 31.77 31.26 31.65 31.72 31.39 31.42 31.50 31.31 31.34 30.90 30.94 30.90 31.01 31.15 30.96 31.04
80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc, 80-551-5cc, 80-551-5cc, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3,
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16^19 16^19 16^19 27^30 39^42 48^51 48^51 48^51 57^60 67^70 78^81 87^90 87^90 87^90 96^99 108^111 108^111 108^111 127^130 139^142 146^150 18^21 18^21 18^21 27^30 37^40 57^60 76^79 76^79 76^79 86^89 106^109 106^109 106^109 116^118 4^7 13^16 13^16 58^61 58^61 58^61 117^120 117^120 117^120 57^60 57^60 57^60 117^120 117^120 117^120 37^40 37^40 37^40
MBS
Planktonic foram.
N13 C
N18 O
132.66 132.66 132.66 132.77 132.89 132.98 132.98 132.98 133.07 133.17 133.18 133.37 133.37 133.37 133.46 133.58 133.58 133.58 133.39 133.51 133.58 134.18 134.18 134.18 134.28 134.38 134.58 134.76 134.76 134.76 134.86 135.06 135.06 135.06 135.16 135.54 135.63 135.63 139.08 139.08 139.08 139.67 139.67 139.67 140.57 140.57 140.57 141.17 141.17 141.17 141.87 141.87 141.87
W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis
2.77 2.61 2.76 2.86 2.92 2.40 2.40 2.47 2.22 2.54 2.53 2.45 2.53 2.47 2.61 2.58 2.40 2.49 2.64 2.65 2.64 2.66 2.55 2.60 2.89 2.39 2.73 2.64 2.65 2.66 2.62 2.60 2.54 2.53 2.47 2.66 2.59 2.62 2.36 2.08 2.16 2.24 2.17 2.20 2.26 2.04 2.09 2.40 2.20 2.23 2.52 2.05 2.27
31.90 32.04 31.76 32.18 33.15 32.94 32.65 32.97 32.65 32.02 32.28 32.37 32.16 32.33 32.48 32.44 32.18 32.29 32.46 32.22 32.39 32.01 31.91 31.94 32.10 33.00 32.16 32.31 32.09 32.46 32.49 32.15 32.32 32.38 32.60 32.68 32.27 32.59 32.05 31.88 31.56 31.62 32.04 31.83 31.75 31.74 31.53 31.58 31.81 31.59 31.86 31.61 31.59
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Table 1 (Continued). MBS
Bulk samples
N13 C
N18 O
Sample
MBS
117^120 138^141 18^21 37^40 60^63 60^63 77^80
141.17 141.38 141.68 141.87 142.10 142.10 142.27
Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment
2.49 2.51 2.39 2.46 2.42 2.43 2.33
31.07 31.26 31.15 31.17 31.51 31.68 -1.24
80-551-6-3, 77^80 80-551-6-3, 77^80 80-551-6-3, 77^80
142.27 H. delrioensis 142.27 R. cushmani 142.27 R. greenhornensis
MBS
Benthic foram.
N13C
N18O
Sample
MBS
6^9 16^19 39^42 39^42 48^51 57^60 67^70 78^81 87^90 96^99 108^111 116^119 127^130 139^142 139^142 146^150 18^21 37^40 76^79 13^16
132.56 132.66 132.89 132.89 132.98 133.07 133.17 133.28 133.37 133.46 133.58 133.66 133.77 133.89 133.89 133.96 134.18 134.37 134.76 135.63
T. T. G. T. G. T. T. T. T. T. T. T. T. T. T. T. T. T. T. T.
1.96 1.87 2.24 2.03 1.69 1.56 1.62 1.74 1.72 1.83 1.85 1.84 1.93 1.93 2.00 1.96 2.06 2.16 2.05 1.95
30.84 31.34 30.59 30.42 31.36 31.17 31.07 31.29 31.18 31.30 31.35 31.18 31.56 31.41 31.47 31.33 31.26 31.16 31.15 31.19
80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3,
138.67 138.87 138.87 139.08 139.32 139.47 139.67 139.89 140.17 140.37 140.57 140.76 140.97 141.17 141.38 141.68 141.87 142.10 142.27 142.27
Sample 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, Sample 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc,
laciniosa laciniosa lenticulus laciniosa lenticulus laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa
37^40 37^40 58^61 82^85 17^20 97^100 117^120 139^141 17^20 37^40 57^60 76^79 97^100 117^120 138^141 18^21 37^40 60^63 77^80 77^80
N13 C
N18 O
2.48 2.11 2.25
32.01 31.94 31.75
Benthic foram.
N13C
N18O
G. T. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. T.
1.42 1.41 1.33 1.56 1.49 1.01 1.24 1.56 1.26 1.15 1.16 1.32 1.53 1.77 1.61 1.38 1.77 1.32 1.50 1.37
30.46 30.77 30.80 30.61 30.60 30.67 30.63 30.65 30.64 30.77 30.68 30.81 30.69 30.56 30.57 30.57 30.57 30.70 30.56 30.92
Planktonic foram.
lenticulus laciniosa lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus laciniosa
All values are expressed as per mil deviations from the PDB standard.
lished for Eastbourne, England (Paul et al., 1999), Drill Hole S75 in the Tarfaya Basin, Morocco (Kuhnt et al., 2001), and ODP Site 1050, Blake Nose (Huber et al., 1999), in the North Atlantic. The locations of these sites and of DSDP Site 551 are plotted on a paleogeographical map for the mid-Cretaceous (92 Ma) in Fig. 1. The main objectives of this study are: (1) To obtain information about the e⁄ciency of the biological pump during di¡erent stages of the CTBI (black shales ^ non-black shales) by comparing surface-, mid- and deep-water carbon isotope records from coccoliths (bulk sediment and ¢ne fraction), shallow dwelling and deeper dwelling planktonic foraminifera and benthic foraminifera. (2) To obtain relative temperature records at di¡erent positions in the water column during
the CTBI from oxygen isotope data of coccoliths, shallow dwelling and deeper dwelling planktonic foraminifera and benthic foraminifera.
2. Materials and methods Site 551 (3909 m water depth) is at the seaward edge of the Goban Spur, southwest of Ireland (48‡54.64PN, 13‡30.09PW). During the late Cretaceous and Paleogene, Site 551 underwent rapid subsidence, resulting into a water depth increase of approximately 1500 m since the CTBI (Masson et al., 1985). The late Cenomanian chalks in Core 551-6R (Fig. 2) are white or pale yellow, and consist of nearly pure calcium carbonate (90^95%, de Graciansky and Bourbon, 1985). A coring gap of almost 3 m occurs below the late Cenomanian
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chalks and the black shale above. The black shale contains 8^11% Total Organic Carbon that is of mainly marine origin (Waples, 1985). Above the black shale, white to pale green massive chalks of latest Cenomanian and early Turonian age occur, which have a generally lower calcium carbonate content (40^74%, de Graciansky and Bourbon, 1985). About 5 g of sediment per sample were treated with 5% hydrogen peroxide at 20‡C. The sediment samples were washed over a 63-Wm sieve, dried at 50‡C, then dry sieved into 63^125-Wm, 125^250Wm and s 250-Wm fractions. Stable isotopes N18 O and N13 C were measured on selected foraminiferal tests, bulk sediment and ¢ne fractions 6 63 Wm. Planktonic and benthic foraminifera are abundant in all samples, except within the black shale, where very few specimens are found. Both calcareous and agglutinated foraminifera are generally well preserved. We used the planktonic foraminiferal species Hedbergella delrioensis, Dicarinella hagni, Rotalipora cushmani, Rotalipora greenhornensis, Whiteinella aprica, and Whiteinella archaeocretacea ( s 250-Wm fraction) and the benthic species Gyroidinoides lenticulus (125^250-Wm fraction) and Tappanina laciniosa (63^125-Wm fraction). The foraminiferal tests were ultrasonically cleaned in ethanol and when possible tests s 125 Wm were crushed before cleaning. Unbroken and broken tests of planktonic foraminifera are illustrated in Fig. 3 to show that the tests are clean and well preserved. To obtain the required s 10 Wg from each sample, about 20 specimens of the larger species D. hagni, H. delrioensis, R. cushmani, R. greenhornensis, W. aprica, W. archaeocretacea, and G. lenticulus and about 50 specimens of the smaller species T. laciniosa were used. The isotope samples were processed in the Carbo Kiel automated carbonate preparation device, linked on-line to a Finnigan MAT 25 mass spectrometer. The analyses were performed at the
57
Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research at Kiel University, Germany. Isotope values are presented in per mil (x) relative to the Pee Dee Belemnite (PDB) standard and NBS-20 standards were used between samples. The analytical precision was better than P 0.08x for N13 C and P 0.08x for N18 O. Isotope measurements are listed in Table 1. The equation from Bemis et al. (1998), T = 16.5^4.80 * (Nc 3Nw ), was used to calculate paleotemperature. Since most authors assume that extensive ice sheets did not exist during the Cretaceous, a N18 Ow of 31.22x estimated for preglacial seawater (Shackleton and Kennett, 1975) was used to estimate the paleotemperature in bottom water. Considering latitudinal di¡erences in surface water N18 Ow during the mid-Cretaceous (Poulsen et al., 1999a), a N18 Ow of 31.47x was used for planktonic foraminifera and bulk sediment. When calculating paleotemperatures, no vital e¡ects were considered.
3. Results 3.1. Biostratigraphy and carbon isotope stratigraphy We correlated the carbonate N13 C isotope curve at Site 551 to the N13 C isotope curve from the Eastbourne section in southern England (Fig. 4), since this is the most expanded and complete CTBI record in NW Europe (Paul et al., 1999). In the late Cenomanian, the Eastbourne carbonate N13 C values show ¢rst a slow increase, then a more rapid increase through the lower half of the Plenus Marls (Paul et al., 1999). The N13 C curve then shows two prominent peaks separated by a trough in the upper half of the Plenus Marls (Paul et al., 1999). The last occurrence of Rotalipora cushmani is within this interval. After the second
Fig. 3. Scanning electron micrographs of broken and unbroken tests of planktonic foraminifera from Site 551. Tests show excellent preservation and little recrystallization. Scale bar equals 100 Wm for specimens 1^5; and 30 Wm for specimen 6. (1) Rotalipora cushmani, Sample 80, 551-6R-2W, 138^141 cm. (2) Rotalipora greenhornensis, Sample 80, 551-6R-2W, 138^141 cm. (3) Dicarinella hagni, Sample 80, 551-5R-1W, 87^90 cm. (4) Dicarinella sp., Sample 80, 551-5R-1W, 87^90 cm. (5) Whiteinella aprica, 80, 551-6R-2W, 138^141 cm. (6) Whiteinella sp., Sample 80, 551-5R-1W, 87^90 cm.
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Fig. 4. Correlation of bulk sediment carbon isotope (N13 C) curves for Tarfaya Drill Hole S75 (Kuhnt et al., 2001), Eastbourne (Paul et al., 1999) and Site 551. Organic carbon was used for the carbon isotope measurements at Tarfaya Drill Hole S75, while carbonate was used for the carbon isotope measurements at the other two sites. The ¢rst occurrence of the nannofossil Quadrum gartneri is indicated in the Eastbourne curve. Rotaliporids are only present in Core 551-5R, their extinction level falls in the gap between Cores 551-5R and -6R.
peak, high carbon isotope values continue into the Holywell Marls, forming a plateau, and ¢nally carbon isotope values decrease gradually (Paul et al., 1999). In Eastbourne, the whole carbon isotope excursion is extended over about 17 m. The ammonite Watinoceras devonense Wright and Kennedy, which is used to de¢ne the base of the Turonian stage (Bengtson, 1996), is not found at Eastbourne. Therefore, the base of the Turonian stage is taken locally between the last occurrence of Neocardioceras juddii and the ¢rst appearance of Fagesia catinus (Gale, 1996). The nannofossil Quadrum gartneri ¢rst occurs at about 0.5 m above the last occurrence of N. juddii, close to the Cenomanian/Turonian boundary proposed by Gale (1996). With respect to the carbon isotope excursion, the Cenomanian/Turonian boundary is placed in the uppermost part of the N13 C plateau at Eastbourne. The presence of Rotalipora cushmani, R. deeckei
and R. greenhornensis and the absence of a carbon isotope excursion throughout Sections 551-6R-1 to -3 indicate a late Cenomanian age for this interval. The lower part of Core 551-5R, which includes the second carbon isotope buildup and peak, is consequently of latest Cenomanian age (Fig. 4). Sections 551-5R-1 to -2 can be correlated with the second carbon isotope peak and decline in the Plenus Marls of Eastbourne, based on the absence of Rotalipora and the presence of a carbon isotope peak and decline in this interval. 3.2. Benthic foraminiferal assemblages The carbonates below and above the black shales at Site 551 contain well preserved planktonic and benthic foraminiferal assemblages. The ratio of planktonic to benthic foraminifera remains extremely high, except within the black shale where both planktonic and benthic forami-
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Fig. 5. Stable isotope plots of bulk sediment, planktonic foraminifera and benthic foraminifera from Site 551. Temperatures are estimated from planktonic foraminiferal and bulk sediment N18 O values. N18 O temperature estimates are 1.2‡C higher for benthic foraminifera, because no correction for latitudinal di¡erences in Cretaceous sea water N18 Ow is applied.
nifera occur rarely. The well diversi¢ed benthic foraminiferal assemblages above and below the black shale commonly include Gavelinella dakotensis, Gyroidinoides lenticulus, Marssonella oxycona, Clavulinoides gaultinus, Praebulimina nannina, Tappanina laciniosa, Osangularia sp., Pleurostomella spp., Spiroplectinata spp. as well as various nodosariids. By contrast, the black shale usually contains only rare specimens of G. dakotensis, G. lenticulus, Osangularia sp. and T. laciniosa. 3.3. Oxygen isotopes in bulk sediment In Sections 551-6R-1 to -3 (138.67^142.30 me-
ter below sea £oor (mbsf)), the N18 O values show marked £uctuations around a mean value of 31.3x (Fig. 5). In Sections 551-5R-1 to -2 (132.56^134.66 mbsf), the N18 O values decrease from 32.2x at 135.66 mbsf to reach a minimum value of 32.9x at 134.76 mbsf, which corresponds to a N13 C peak. The N18 O values then generally increase above the black shale reaching 31.5xat 132.56 mbsf. Light and SEM microscope observations indicate that the bulk sediment generally consists of coccoliths (Fig. 6). To test whether the bulk signal was a¡ected by other components (such as fragments of foraminiferal tests), we analyzed parallel
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samples of ¢ne fractions 6 63 Wm. The di¡erence between bulk sediment and ¢ne fractions 6 63 Wm was found to be small. For a set of nine parallel samples, the mean di¡erence is 0.06x for N18 O and 0.14x for N13 C. Thus, it can be concluded that the isotope signal in the bulk sediment mainly represents the 6 63-Wm fraction (principally composed of coccoliths). 3.4. Oxygen isotopes in foraminifera In Core 551-6R, the planktonic foraminifera Hedbergella delrioensis show N18 O values varying between 31.6 and 32.05x with a mean value of 31.8x. Values for Rotalipora cushmani and R. greenhornensis are generally about 0.3x higher, and bulk sediment 0.3^0.8x higher than H. delrioensis (Fig. 5). In Core 551-5R, the N18 O values of Whiteinella aprica ¢rst decrease within the black shale, then show an overall increase, whereas the N18 O values for bulk sediment exhibit an overall decline. Above the black shale, Dicarinella hagni, W. aprica and Whiteinella archaeocretacea display N18 O values that are generally similar to the N18 O values in bulk sediment. In the upper part of the section, the W. aprica N18 O curve displays a minimum of 33.2x at 132.89 mbsf. The same pattern is evident in the N18 O curves of D. hagni, W. archaeocretacea and in the bulk sediment curve. In Core 551-6R, the benthic foraminifera Gyroidinoides lenticulus displays stable N18 O values around a mean of 30.7x (Fig. 5). In Core 551-5R, the Tappanina laciniosa record shows a small N18 O peak (31.6x) at 133.77 mbsf. 3.5. Carbon isotopes in bulk sediment In Sections 551-6R-1 to -3 (138.67^142.30 mbsf), the bulk sediment shows persistent N13 C values around a mean of 2.4x (Figs. 3 and 5). In Sections 551-5R-1 and -2 (132.56^135.66 mbsf), the N13 C values in bulk sediment increase from 3.3x at 135.66 mbsf through the black shale and peak to 3.9x at 134.76 mbsf. Above the black shale, the N13 C values decrease from 3.1x at 134.37 mbsf to 2.1x at 132.56 mbsf, except for a smaller peak (3.2x) at 132.89 mbsf.
Fig. 6. Scanning electron micrographs of bulk sediment fraction 6 63 Wm, showing rich coccolith content. The upper micrograph is from Sample 80-551-5R-1, 16^19 cm above the black shale, the middle micrograph is from Sample 80-5515R-cc, 4^7cm, in the uppermost part of the black shale and the lower micrograph is from Sample 80-551-6R-3, 60^63 cm below the black shale.
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As for oxygen isotopes, the di¡erence between bulk sediment and ¢ne fractions 6 63 Wm was found to be small (mean di¡erence of 0.14x for nine parallel measurements). 3.6. Carbon isotopes in foraminifera In Core 551-6R, the planktonic foraminifera Hedbergella delrioensis shows stable N13 C values around a mean value of 2.4x. Rotalipora cushmani, R. greenhornensis and bulk sediment samples show similar values. In Core 551-5R, the N13 C values for Whiteinella aprica, Dicarinella hagni and Whiteinella archaeocretacea are generally stable around a mean value of 2.6x. These planktonic foraminiferal N13 C values are approximately 1x lower than bulk sediment values within the black shale (Fig. 5). In Core 551-6R, the benthic foraminiferal N13 C values are more variable around a mean of 1.44x. The N13 C values in Tappanina laciniosa are relatively stable around 2x within the black shale. Above the black shale, the N13 C values in T. laciniosa decrease by about 0.5x. In the interval 551-5R-1, 6^9 cm to 551-5R-2, 48^50 cm (132.56^134.48 mbsf), the N13 C values in T. laciniosa are overall at least 0.5x lower than for planktonic foraminifera and bulk sediment. Within the black shale (134.76^135.63 cm), the di¡erence between benthic foraminiferal and bulk sediment N13 C values increases to 1.5x.
4. Discussion 4.1. The paleotemperature record of Site 551 and its importance for the reconstruction of the Cenomanian/Turonian Greenhouse climate 4.1.1. Sea surface temperature at Site 551 during the CTBI We reconstructed Cenomanian/Turonian Sea surface temperature (SST) values from three different sets of N18 O measurements: bulk sediment, Hedbergella delrioensis tests and a combination of Whiteinella aprica and Whiteinella archaeocretacea tests, as the latter two species appear to have had very similar habitat and isotope frac-
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tionation. Bulk sediment and 6 63-Wm fractions consist almost exclusively of coccoliths, thus their oxygen isotope values re£ect temperatures near the sea surface. However, the reliability of these data is limited by: (1) diagenetic alteration, (2) unknown vital e¡ects in oxygen isotope fractionation of Cretaceous coccolithophorids, and (3) extreme seasonality of coccolithophorids which commonly represent the population of a single plankton bloom with a duration of a few days. The planktonic foraminiferal genera Hedbergella and Whiteinella are commonly assumed to have lived predominantly at shallow depths in surface waters. This assumption is based on morphologic comparison with modern planktonic foraminiferal taxa (Hart, 1980), their distribution in marginal areas (Leckie, 1987) and stable isotope values (D’Hondt and Arthur, 1995; Huber et al., 1995). Whiteinella aprica exhibits N18 O values in Core 551-5R that are similar to those of bulk sediment. Thus, we regard paleotemperature estimates of 18.6^23.8‡C for W. aprica and 16.6^23.4‡C for bulk sediment as reasonably reliable for surface water. In contrast, Hedbergella delrioensis displays consistently lower N18 O values than coccolith bulk sediment in Core 551-6R. Thus, paleotemperature estimates for H. delrioensis (17.0^19.3‡C) are generally higher than for bulk sediment (13.8^ 21.4‡C). This o¡set in N18 O values between late Cenomanian surface dwelling planktonic foraminifera and coccoliths (bulk sediment) may be explained by the existence of a strong seasonal contrast in the late Cenomanian. Coccoliths may have been produced during the colder spring while planktonic foraminifera grew during the warmer summer. The similarity in N18 O values between planktonic foraminifera and bulk sediment in the latest Cenomanian^early Turonian (Core 551-5R) may indicate less pronounced seasonal temperature variations in surface water. Alternative explanations would be di¡erent vital e¡ects for W. aprica and H. delrioensis or signi¢cantly di¡erent composition of coccolith assemblages in Cores 551-5R and -6R. A clear warming trend from the latest Cenomanian to the earliest Turonian is observed. Average SST estimates, based on planktonic foraminiferal N18 O values corrected for latitudinal di¡erence in
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surface water N18 O (Poulsen et al., 1999a), rise from 17^19‡C in Core 551-6R (late Cenomanian) to a maximum of 23‡C in the latest Cenomanian and earliest Turonian, within and directly above the black shale. These results are in accordance with the study by Frakes (1999), which indicated a colder climate during the middle/late Cenomanian (Core 551-6R) and increasing temperatures with a temperature maximum in the latest Cenomanian and early Turonian (Core 551-5R). Using pristine planktonic foraminifera, Pearson et al. (2001) estimated warmer tropical surface water paleotemperatures for the Maastrichtian and mid-Eocene than had been estimated before. This ¢nding highlighted the importance of test preservation and the in£uence that secondary calcite had on paleotemperature estimates (Pearson et al., 2001). Even for our well preserved material from Site 551, a minor contribution from secondary calcite crystals cannot be excluded. Thus the absolute values for our temperature estimates should be treated with caution. 4.1.2. Comparison with records from other localities In the Eastbourne C/T section, situated at comparable paleolatitude as Site 551, most bulk sediment N18 O values are about 1x lower than at Site 551 (Paul et al., 1999). However, the sediments from Eastbourne experienced signi¢cant diagenesis and the N18 O values for chalk are approximately 1.0^1.5x more negative than those for adjacent marls, which were less in£uenced by diagenetic processes (Paul et al., 1999). When this diagenetic alteration is taken into account, the early Turonian SST estimates at Eastbourne are within the range of those estimated for Site 551. More reliable and less diagenetically altered is the Cenomanian/Turonian paleotemperature record based on oxygen isotopes of planktonic foraminifera from ODP Site 1050 at Blake Nose, at approximately 27‡N paleolatitude. However, the sedimentary record of the Cenomanian/Turonian transition is not complete at Site 1050. The second buildup and peak in the carbon isotope curve, documented at Eastbourne (Paul et al., 1999) and at Site 551 (Fig. 4), are not present in the Blake Nose record (Huber et al., 1999). Dur-
ing the recovery after the second carbon isotope buildup and plateau, planktonic foraminifera show similar N18 O and N13 C values at both Sites 1050 and 551. In the subtropical (15‡N paleolatitude) eastern North Atlantic, N18 O paleotemperatures between 30 and 35‡C were estimated in the Tarfaya coastal basin, southern Morocco for the interval representing the CTBI carbon isotope excursion (Luderer and Kuhnt, 1997). These temperatures may be overestimated due to diagenetic alteration, and are approximately 5‡C too warm when compared to Poulsen et al.’s (1999a) Parallel Ocean Climate Model simulations. However, well preserved Turonian planktonic foraminifera from the western tropical Atlantic also yielded estimated N18 O temperatures of at least 30^33‡C (Wilson et al., 2002). A late Cenomanian warming trend and a temperature maximum at approximately the extinction level of Rotalipora cushmani followed by stepwise cooling, which are evident in the Tarfaya record, are also found in the Blake Nose, Eastbourne and Goban Spur records. 4.1.3. Bottom-water temperatures Brass et al. (1982) proposed that evaporation in subtropical marginal seas and the subsequent sinking of warm and saline waters were important processes for the formation of deep water in the Cretaceous. Models for mid-Cretaceous ocean circulation indicate strong downward convection of warm and saline waters from subtropical areas in the eastern Tethys (Barron and Peterson, 1990). High density, warm and saline water may have also been generated in the Gulf of Mexico and contributed to intermediate and deep-water formation (Woo et al., 1992). Increases in bottomwater temperatures strongly in£uence the development of low oxygen conditions in the deep sea in two main ways: (1) it leads to increased metabolism in exothermic organisms, and (2) decreases the ability for water to dissolve oxygen. However, it is still a matter of vigorous discussion whether the presence of warm, anoxic, saline deep water alone without signi¢cantly increased surface productivity could cause deep-water anoxia and enhance the preservation of black shales in the deep sea.
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Deep-water paleotemperature estimates are more reliable since benthic foraminifera seem less prone to diagenetic alteration than planktonic foraminifera (Pearson et al., 2001). Thus, paleotemperature estimates of 12.9^16.9‡C (Gyroidinoides lenticulus) for Core 551-6R and 12.7^ 18.1‡C (Tappanina laciniosa) for Core 551-5R are considered as reasonably reliable for bottom water. Benthic foraminiferal N18 O values at Site 551 indicate no signi¢cant change in deep-water temperatures at the end of black shale deposition. Relatively uniform bottom-water temperatures through and immediately above the black shale (Fig. 5) suggest that elevated bottom-water temperatures were not the only reason for the development of low oxygen conditions at Site 551. It is likely that increased evaporation leading to the production of warm, highly saline and dense water during the CTBI occurred in Tethyan subtropical seas. On its way down, the warm and dense water plume may have entrained intermediate water and gradually become less saline and colder, before spreading out laterally as the Mediterranean water spreads into the Atlantic today. Thus, the warm climate during the CTBI may have stimulated the production of intermediate and deep water, in turn increasing vertical turnover and upwelling, and leading to increased productivity. If this scenario holds, the very warm climate at the CTBI could, at least partly, explain the increase in productivity and the deposition of organic rich sediments during this event. The warmest deep-water temperatures during the Cretaceous are recorded in the latest Cenomanian and early Turonian, when temperatures averaged 19‡C (Huber et al., 2002). At Site 551 (Core 6), Cenomanian deep-water paleotemperatures of approximately 13 to 16‡C are similar to or slightly lower than at Site 1050. Maximum temperatures of approximately 17‡C at Site 551 and 18‡C at Site 1050 follow the same trend during the isotope excursion. After the second carbon isotope peak bottom water, temperatures decreased from approximately 18 to 17‡C at Site 1050 (Huber et al., 1995) and from approximately 17 to 15‡C at Site 551. The slightly lower paleotemperatures at Site 551 are easily explained by the shallower pa-
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leodepth at Site 1050 (around 1200 m) than at Site 551 (around 2400 m). 4.2. Benthic foraminiferal distribution during the low oxygen event Within the black shale in Core 551-5R, the benthic foraminiferal assemblages display low diversity in contrast to intervals above and below the black shale. This is similar to the distribution in modern anoxic or close to anoxic marine sediments, where benthic foraminiferal assemblages generally show low diversity and low abundance (Bernhard et al., 1997; Gustafsson and Nordberg, 2000). Species within the black shale include Osangularia sp., Gyroidinoides lenticulus, Gavelinella dakotensis and Tappanina laciniosa, which are characteristic species of Albian and late Cenomanian black shales (Koutsoukos et al., 1990; Tronchetti and Grosheny, 1991; Kuhnt and Wiedmann, 1995; Erbacher et al., 1998; Holbourn et al., 2001). T. laciniosa and G. dakotensis, in particular, are characterisitc elements of late Cenomanian Tethyan and temperate shelf biofacies with enhanced carbon £ux and dysoxic bottom waters (Kuhnt and Wiedmann, 1995; Holbourn and Kuhnt, 2002). Site 551 represents the deepest occurrence of this typical shelf/slope high productivity assemblage (2400 m paleowater depth). This probably indicates an unusually high carbon £ux at this depth that only can be explained by considerably enhanced primary production. 4.3. Carbon £ux during the CTBI We observe a signi¢cant deviation between the N13 C curves of surface dwelling coccoliths (bulk sediment) and deeper dwelling non-keeled planktonic foraminifera within the CTBI black shale (Fig. 5). Above and below the black shale, planktonic foraminifera (Whiteinella aprica in the upper part of Core 551-5R and Hedbergella delrioensis in Core 551-6R) exhibit N13 C values similar to that of coccolith bulk sediment. However, within the black shale W. aprica and W. archaeocretacea have signi¢cantly lower N13 C values than bulk sediment, and show a minimum in N13 C values
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or no N13 C peak. A steep N13 C gradient through the upper water column during black shale deposition would best account for the deviation in N13 C signals, which can only be caused by signi¢cantly increased primary production and carbon £ux rates. Since both coccoliths and planktonic foraminifera live in the photic zone, the N13 C gradient must have extended from the water surface down to a few tenths of meters. The small depth extent of the N13 C gradient is also supported by the similarity in the oxygen isotope values of coccoliths and planktonic foraminifera. The similarity in the shape of the N13 C curves for coccoliths and benthic foraminifera suggests that the N13 C excursion was caused by global reservoir changes. However, the benthic N13 C excursion has a signi¢cantly lower amplitude (generally less than 1x) than the coccolith curve, which can be explained by the fact that the remineralization of large amounts of organic material signi¢cantly decreased N13 C values in sediment pore waters. Wilson and Norris (2001) proposed that a breakdown in water strati¢cation during the late Albian (OAE 1d) led to increased upwelling, increased production and subsequently to low oxygen bottom conditions and deposition of organic rich sediments. The situation at the CTBI may have been comparable. Weak water temperature strati¢cation would mean increased upmixing/upwelling of nutrient rich deepwater. This would lead to increased production and to the development of a steep N13 C surface water gradient, as well as increased export of organic material and subsequent development of anoxic or dysoxic conditions at the sea£oor during black shale deposition. Within the black shale at Site 551, we observe an unusual steep carbon isotope gradient between surface and bottom water, which can be accounted for by high productivity and an extremely active biological pump. High productivity is not possible in strongly strati¢ed water masses, since the uptransport of nutrients through the pycnocline is very limited. As Site 551 is far from the coastline, a signi¢cant nutrient supply from land is also not likely. At Sites 1050 and 551, planktonic and benthic foraminiferal N13 C values through the carbon isotope excursion are
comparable in magnitude. However, at Site 1050 planktonic foraminifera show a distinct N13 C peak through the black shale (Huber et al., 1999), which is not the case at Site 551 as mentioned earlier. It is therefore likely that the surface water N13 C gradient was not so steep at Site 1050, and primary production was possibly lower than at Site 551.
5. Conclusions The good preservation of calcareous microfossils together with the great paleodepth (about 2400 m) of DSDP Site 551 at Goban Spur (midlatitude NE Atlantic) provide a unique opportunity to use stable oxygen and carbon isotope measurements to reconstruct vertical temperature gradients and to assess the e⁄ciency of the biological pump in the water column during the CTBI. Stable isotope data of bulk sediment (coccolith) and well preserved tests of planktonic and benthic foraminifera allow to quantify surface and deepwater temperature changes that occurred during the CTBI. SSTs of about 18‡C and deep-water temperatures of about 14‡C indicate a low temperature gradient within the water column during the late Cenomanian. Following additional warming of at least 2‡C in both surface and deep water in the latest Cenomanian reaching a temperature maximum during black shale deposition, temperatures decreased again slightly in the early Turonian. Even in our well preserved material, a minor contribution of secondary calcite crystals cannot be excluded. Although absolute temperature values should be treated with some caution, temperature gradients are generally not in£uenced by this process. The CTBI carbon isotope excursion is prominent in the bulk (coccolith) signal, but less pronounced in the planktonic and benthic foraminiferal records. This di¡erence indicates a very steep 13 C gradient in the upper water column within the photic zone and a very e⁄cient biological pump during the CTBI carbon isotope excursion. We suggest signi¢cantly increased primary production in the uppermost water column and a shallow chlorophyll maximum within a water depth of
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few meters or tens of meters as the cause for this steep carbon isotope gradient. The temperature changes at Goban Spur exhibit the same trend of warming towards the CTBI black shale interval as at low latitude locations: ODP Site 1050 Blake Nose, and Tarfaya, Morocco. The latitudinal temperature gradient is in agreement with the (GCM) simulations by Poulsen et al. (1999a, b).
Acknowledgements We thank Dr. A. Niederbragt and Professor I. Premoli Silva for very helpful reviews. We also extend our thanks to Dr. Helmut Erlenkeuser at the Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research, to the sta¡ of the SEM Laboratory at the Institut fu«r Geowissenschaften, both at the Christian Albrechts University in Kiel and to the sta¡ of the ODP Core Repository in Bremen. This research used samples provided by the Ocean Drilling Program. The ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions Inc. We gratefully acknowledge ¢nancial support from the E.U. ¢nanced program ‘Rapid global change during the Cenomanian/Turonian OAE : Examination of a natural climatic experiment in earth history’, Grant HPRN-CT 199900055. The Deutsche Forschungsgemeinschaft is also gratefully acknowledged for ¢nancial support (Grants KU 649/7 and KU 649/8).
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Geology of the Goban Spur Continental Margin. Initial Reports of the Deep Sea Drilling Project 80, U.S. Government Printing O⁄ce, Washington, DC, pp. 1115^1151. Paul, C.R.C., Lamolda, M.A., Mitchell, S.F., Vaziri, M.R., Gorostidi, A., Marshall, J.D., 1999. The Cenomanian^Turonian boundary at Eastbourne (Sussex, UK): A proposed European reference section. Paleogeogr. Paleoclimatol. Paleoecol. 150, 83^121. Pearson, P.N., Ditch¢eld, P.W., Singano, J., Harcourt-Brown, K.G., Nicholas, C.J., Olsson, R.K., Shackleton, N.J., Hall, M.A., 2001. Warm tropical sea surface temperatures in the Late Cretaceous and Eocene epochs. Nature 413, 481^487. Poulsen, C.J., Barron, E.J., Peterson, W.H., Wilson, P.A., 1999a. A reinterpretation of mid-Cretaceous shallow marine temperatures through model-data comparison. Paleoceanography 14, 679^697. Poulsen, C.J., Barron, E.J., Johnson, C.C., Fawcett, P., 1999b. Links between major climatic factors and regional oceanic circulation in the mid-Cretaceous. In: Barrera, E., Johnson, C.C. (Eds.), Evolution of the Cretaceous Ocean-Climate System. Geological Society of America Spec. Publ. 332, pp. 73^89. Shackleton, N.J., Kennett, J.P., 1975. Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: Oxygen and carbon isotope analysis in DSDP Sites 277, 279 and 280. In: Kennett, J.P., Houtz, R.E., and others, Initial Reports of the Deep Sea Drilling Project 29, U.S. Government Printing O⁄ce, Washington, DC, pp. 743^755. Tronchetti, G., Grosheny, D., 1991. Les assemblages de foraminife'res benthiques au passage Ce¤nomanien^Turonien a¤ Vergons, S-E France. Geobios 24, 13^31. Waples, D.W., 1985. A reappraisal of anoxia and richness of organic material, with emphasis on the Cretaceous north Atlantic. Initial reports of the Deep Sea Drilling Project 80, part 2, U.S. Government Printing O⁄ce, Washington, DC, pp. 999^1016 Wilson, A.W., Norris, R.D., 2001. Warm tropical ocean surface and global anoxia during the mid-Cretaceous period. Nature 412, 425^429. Wilson, A.W., Norris, R.D., Cooper, M.J., 2002. Testing the Cretaceous greenhouse hypothesis using glassy foraminiferal calcite from the core of the Turonian tropics on Demerara Rise. Geol. Soc. Am. 30, 607^610. Woo, K-S., Anderson, T.F., Railback, L.B., Sandberg, P.A., 1992. Oxygen isotope evidence for high salinity surface seawater in the mid-Cretaceous Gulf of Mexico: Implications for warm, saline deepwater formation. Paleoceanography 7, 673^685.
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Changes in Northeast Atlantic temperature and carbon £ux during the Cenomanian/Turonian paleoceanographic event: the Goban Spur stable isotope record M. Gustafsson , A. Holbourn, W. Kuhnt Institut fu«r Geowissenschaften, Christian-Albrechts-Universita«t, Olshausenstrasse 40, D-24118 Kiel, Germany Received 20 November 2001; received in revised form 2 June 2003; accepted 17 June 2003
Abstract Stable isotopes of bulk sediment and well preserved tests of planktonic and benthic foraminifera from midlatitude NE Atlantic DSDP Site 551 (Goban Spur) provide the first estimates of carbon isotope gradients within the water column at a lower bathyal site during the Cenomanian/Turonian boundary interval (CTBI). The CTBI carbon isotope excursion is prominent (up to 2x shift in N13 C) in the bulk (coccolith) signal, but less pronounced (approximately 0.5x shift in N13 C) in planktonic and benthic foraminifera. This difference indicates a very steep 13 C gradient in the upper water column and a very efficient biological pump during the CTBI carbon isotope excursion. We suggest significantly increased seasonal primary production in the uppermost water column with an enhanced shallow water chlorophyll maximum as a cause for this steep carbon isotope gradient. Deep-water and surface-water temperature changes during the CTBI are estimated using benthic and planktonic foraminiferal oxygen isotopes. Warm deep-water masses (13^16‡C) and a low temperature gradient within the water column prevailed in the late Cenomanian. Additional warming (approximately 2‡C for both surface and deep water) occurred in the latest Cenomanian prior to CTBI black shale deposition. This pattern of CTBI black shale deposition during a temperature maximum is also evident at two low latitude locations (ODP Site 1050, Blake Nose and Tarfaya, southern Morocco). < 2003 Published by Elsevier B.V. Keywords: Cenomanian^Turonian boundary; stable isotopes; foraminifer; carbon cycle; NE Atlantic; paleotemperature
1. Introduction The mid-Cretaceous represents one of the most remarkable episodes of greenhouse climate in Earth’s history. It was characterized by high atmospheric CO2 concentrations, low latitudinal temperature gradient and unusually high deep-
* Corresponding author. Fax: +49-431-880-4376. E-mail address: [email protected] (M. Gustafsson).
water temperatures in the global ocean (Hay, 1995). The mid-Cretaceous greenhouse was coincident with a worldwide pulse in ocean crustal production (Larson, 1991). The release of mantle CO2 from this very active volcanic episode may have in fact directly caused the warm mid-Cretaceous greenhouse climate (Larson, 1991). Thus, the study of mid-Cretaceous paleoclimate and paleoceanography provides insight into a natural climatic experiment, when a large amount of CO2 was released into the atmosphere. During
0031-0182 / 03 / $ ^ see front matter < 2003 Published by Elsevier B.V. doi:10.1016/S0031-0182(03)00509-1
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Fig. 1. Paleogeographical map of the mid-Cretaceous (92 Ma) North Atlantic showing the position of DSDP Site 551, ODP Site 1050, Drill Hole S75 and the Eastbourne section. Present day shorelines are indicated and gray shaded areas represent continental reconstruction from data ¢les in Hay et al. (1999), downloaded from http://www.odsn.de/odsn/services/paleomap.
the mid-Cretaceous, widespread deposition of organic carbon occurred during several oceanic anoxic events (OAEs). Erbacher et al. (1999) de¢ned ¢ve OAEs, and the Cenomanian/Turonian boundary interval (CTBI) represents one of the more severe and long-lasting of those events. Caron et al. (1999) estimated the duration of the CTBI as 0.4 myr., but both longer and shorter durations have been proposed by other authors. A positive carbon isotope excursion of up to 2x in carbonate carbon and up to 6x in organic carbon is associated with the CTBI (Arthur et al., 1988). Most likely, the carbon isotope excursion resulted from the removal of light 12 C from the global reservoir by the globally enhanced burial of 13 Cdepleted organic carbon into black shales (Arthur et al., 1988). However, it is still a matter of dis-
cussion whether this increased burial of organic carbon was mainly caused by high primary production (e⁄cient biological pump) or by enhanced organic matter preservation at comparatively low primary production levels. We present new oxygen and carbon isotope data across the CTBI from Deep Sea Drilling Program (DSDP) Site 551, which is located on the Goban Spur in the eastern North Atlantic (de Graciansky and Bourbon, 1985). The unusually good preservation of calcareous foraminifera and a paleodepth close to, but above the Calcite Compensation Depth a¡ord a unique opportunity to estimate both the carbon isotope gradient within the water column and the oxygen isotopic paleotemperatures of deep and surface waters. The isotopic results are compared with those pub-
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Fig. 2. Bulk sediment carbon isotope (N13 C) curve and lithological summary for Site 551 (modi¢ed after de Graciansky and Bourbon, 1985).
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Table 1 Oxygen and carbon isotope data for bulk sediment, fraction 6 63 Wm, planktonic foraminifera and benthic foraminifera from Site 551. Sample 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc, 80-551-5cc, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2,
6^9 16^19 27^30 39^42 39^42 48^51 67^70 67^70 78^81 87^90 96^99 96^99 108^111 116^119 127^130 127^130 139^142 139^142 146^150 6^9 18^21 18^21 18^21 27^30 37^40 48^50 57^60 76^79 86^89 96^99 106^109 106^109 116^118 4^7 13^16 17^20 37^40 37^40 58^61 82^85 97^100 117^120 117^120 139^141 17^20 17^20 37^40 57^60 57^60 76^79 76^79 76^79 97^100
MBS
Bulk samples
N13 C
N18 O
Sample
132.56 132.66 132.77 132.89 132.89 132.98 133.07 133.07 133.28 133.37 133.46 133.46 133.58 133.66 133.77 133.77 133.89 133.89 133.96 134.06 134.18 134.18 134.18 134.27 134.37 134.48 134.57 134.76 134.86 134.96 135.06 135.06 135.16 135.54 135.63 138.67 138.87 138.87 139.08 139.32 139.47 139.67 139.67 139.89 140.17 140.17 140.37 140.57 140.57 140.76 140.76 140.76 140.97
Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment
2.12 2.77 2.84 3.18 3.12 2.84 2.64 2.63 2.49 2.73 2.64 2.66 2.64 2.75 2.66 2.66 2.70 2.69 2.69 3.18 2.90 2.89 2.93 2.94 3.05 3.18 3.80 3.92 3.68 3.52 3.01 3.01 3.67 3.43 3.28 2.28 2.43 2.57 2.33 2.58 2.43 2.33 2.43 2.35 2.29 2.35 2.20 2.25 2.28 2.16 2.09 2.22 2.40
31.49 31.88 31.89 31.63 31.52 32.56 32.08 31.93 32.27 32.09 32.36 32.11 32.12 31.90 32.33 32.28 32.04 32.07 32.14 32.63 31.88 32.09 31.90 32.16 31.79 31.89 32.57 32.92 32.59 32.57 32.29 32.20 32.50 32.46 32.22 31.39 31.38 31.77 31.26 31.65 31.72 31.39 31.42 31.50 31.31 31.34 30.90 30.94 30.90 31.01 31.15 30.96 31.04
80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc, 80-551-5cc, 80-551-5cc, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3,
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16^19 16^19 16^19 27^30 39^42 48^51 48^51 48^51 57^60 67^70 78^81 87^90 87^90 87^90 96^99 108^111 108^111 108^111 127^130 139^142 146^150 18^21 18^21 18^21 27^30 37^40 57^60 76^79 76^79 76^79 86^89 106^109 106^109 106^109 116^118 4^7 13^16 13^16 58^61 58^61 58^61 117^120 117^120 117^120 57^60 57^60 57^60 117^120 117^120 117^120 37^40 37^40 37^40
MBS
Planktonic foram.
N13 C
N18 O
132.66 132.66 132.66 132.77 132.89 132.98 132.98 132.98 133.07 133.17 133.18 133.37 133.37 133.37 133.46 133.58 133.58 133.58 133.39 133.51 133.58 134.18 134.18 134.18 134.28 134.38 134.58 134.76 134.76 134.76 134.86 135.06 135.06 135.06 135.16 135.54 135.63 135.63 139.08 139.08 139.08 139.67 139.67 139.67 140.57 140.57 140.57 141.17 141.17 141.17 141.87 141.87 141.87
W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. archaeocretacea D. hagni W. aprica W. aprica W. aprica W. aprica H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis H. delrioensis R. cushmani R. greenhornensis
2.77 2.61 2.76 2.86 2.92 2.40 2.40 2.47 2.22 2.54 2.53 2.45 2.53 2.47 2.61 2.58 2.40 2.49 2.64 2.65 2.64 2.66 2.55 2.60 2.89 2.39 2.73 2.64 2.65 2.66 2.62 2.60 2.54 2.53 2.47 2.66 2.59 2.62 2.36 2.08 2.16 2.24 2.17 2.20 2.26 2.04 2.09 2.40 2.20 2.23 2.52 2.05 2.27
31.90 32.04 31.76 32.18 33.15 32.94 32.65 32.97 32.65 32.02 32.28 32.37 32.16 32.33 32.48 32.44 32.18 32.29 32.46 32.22 32.39 32.01 31.91 31.94 32.10 33.00 32.16 32.31 32.09 32.46 32.49 32.15 32.32 32.38 32.60 32.68 32.27 32.59 32.05 31.88 31.56 31.62 32.04 31.83 31.75 31.74 31.53 31.58 31.81 31.59 31.86 31.61 31.59
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55
Table 1 (Continued). MBS
Bulk samples
N13 C
N18 O
Sample
MBS
117^120 138^141 18^21 37^40 60^63 60^63 77^80
141.17 141.38 141.68 141.87 142.10 142.10 142.27
Bulk sediment Bulk sediment Bulk sediment Bulk sediment Bulk sediment 6 63 Wm fraction Bulk sediment
2.49 2.51 2.39 2.46 2.42 2.43 2.33
31.07 31.26 31.15 31.17 31.51 31.68 -1.24
80-551-6-3, 77^80 80-551-6-3, 77^80 80-551-6-3, 77^80
142.27 H. delrioensis 142.27 R. cushmani 142.27 R. greenhornensis
MBS
Benthic foram.
N13C
N18O
Sample
MBS
6^9 16^19 39^42 39^42 48^51 57^60 67^70 78^81 87^90 96^99 108^111 116^119 127^130 139^142 139^142 146^150 18^21 37^40 76^79 13^16
132.56 132.66 132.89 132.89 132.98 133.07 133.17 133.28 133.37 133.46 133.58 133.66 133.77 133.89 133.89 133.96 134.18 134.37 134.76 135.63
T. T. G. T. G. T. T. T. T. T. T. T. T. T. T. T. T. T. T. T.
1.96 1.87 2.24 2.03 1.69 1.56 1.62 1.74 1.72 1.83 1.85 1.84 1.93 1.93 2.00 1.96 2.06 2.16 2.05 1.95
30.84 31.34 30.59 30.42 31.36 31.17 31.07 31.29 31.18 31.30 31.35 31.18 31.56 31.41 31.47 31.33 31.26 31.16 31.15 31.19
80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-1, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3,
138.67 138.87 138.87 139.08 139.32 139.47 139.67 139.89 140.17 140.37 140.57 140.76 140.97 141.17 141.38 141.68 141.87 142.10 142.27 142.27
Sample 80-551-6-2, 80-551-6-2, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, 80-551-6-3, Sample 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-1, 80-551-5-2, 80-551-5-2, 80-551-5-2, 80-551-5cc,
laciniosa laciniosa lenticulus laciniosa lenticulus laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa laciniosa
37^40 37^40 58^61 82^85 17^20 97^100 117^120 139^141 17^20 37^40 57^60 76^79 97^100 117^120 138^141 18^21 37^40 60^63 77^80 77^80
N13 C
N18 O
2.48 2.11 2.25
32.01 31.94 31.75
Benthic foram.
N13C
N18O
G. T. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. G. T.
1.42 1.41 1.33 1.56 1.49 1.01 1.24 1.56 1.26 1.15 1.16 1.32 1.53 1.77 1.61 1.38 1.77 1.32 1.50 1.37
30.46 30.77 30.80 30.61 30.60 30.67 30.63 30.65 30.64 30.77 30.68 30.81 30.69 30.56 30.57 30.57 30.57 30.70 30.56 30.92
Planktonic foram.
lenticulus laciniosa lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus lenticulus laciniosa
All values are expressed as per mil deviations from the PDB standard.
lished for Eastbourne, England (Paul et al., 1999), Drill Hole S75 in the Tarfaya Basin, Morocco (Kuhnt et al., 2001), and ODP Site 1050, Blake Nose (Huber et al., 1999), in the North Atlantic. The locations of these sites and of DSDP Site 551 are plotted on a paleogeographical map for the mid-Cretaceous (92 Ma) in Fig. 1. The main objectives of this study are: (1) To obtain information about the e⁄ciency of the biological pump during di¡erent stages of the CTBI (black shales ^ non-black shales) by comparing surface-, mid- and deep-water carbon isotope records from coccoliths (bulk sediment and ¢ne fraction), shallow dwelling and deeper dwelling planktonic foraminifera and benthic foraminifera. (2) To obtain relative temperature records at di¡erent positions in the water column during
the CTBI from oxygen isotope data of coccoliths, shallow dwelling and deeper dwelling planktonic foraminifera and benthic foraminifera.
2. Materials and methods Site 551 (3909 m water depth) is at the seaward edge of the Goban Spur, southwest of Ireland (48‡54.64PN, 13‡30.09PW). During the late Cretaceous and Paleogene, Site 551 underwent rapid subsidence, resulting into a water depth increase of approximately 1500 m since the CTBI (Masson et al., 1985). The late Cenomanian chalks in Core 551-6R (Fig. 2) are white or pale yellow, and consist of nearly pure calcium carbonate (90^95%, de Graciansky and Bourbon, 1985). A coring gap of almost 3 m occurs below the late Cenomanian
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chalks and the black shale above. The black shale contains 8^11% Total Organic Carbon that is of mainly marine origin (Waples, 1985). Above the black shale, white to pale green massive chalks of latest Cenomanian and early Turonian age occur, which have a generally lower calcium carbonate content (40^74%, de Graciansky and Bourbon, 1985). About 5 g of sediment per sample were treated with 5% hydrogen peroxide at 20‡C. The sediment samples were washed over a 63-Wm sieve, dried at 50‡C, then dry sieved into 63^125-Wm, 125^250Wm and s 250-Wm fractions. Stable isotopes N18 O and N13 C were measured on selected foraminiferal tests, bulk sediment and ¢ne fractions 6 63 Wm. Planktonic and benthic foraminifera are abundant in all samples, except within the black shale, where very few specimens are found. Both calcareous and agglutinated foraminifera are generally well preserved. We used the planktonic foraminiferal species Hedbergella delrioensis, Dicarinella hagni, Rotalipora cushmani, Rotalipora greenhornensis, Whiteinella aprica, and Whiteinella archaeocretacea ( s 250-Wm fraction) and the benthic species Gyroidinoides lenticulus (125^250-Wm fraction) and Tappanina laciniosa (63^125-Wm fraction). The foraminiferal tests were ultrasonically cleaned in ethanol and when possible tests s 125 Wm were crushed before cleaning. Unbroken and broken tests of planktonic foraminifera are illustrated in Fig. 3 to show that the tests are clean and well preserved. To obtain the required s 10 Wg from each sample, about 20 specimens of the larger species D. hagni, H. delrioensis, R. cushmani, R. greenhornensis, W. aprica, W. archaeocretacea, and G. lenticulus and about 50 specimens of the smaller species T. laciniosa were used. The isotope samples were processed in the Carbo Kiel automated carbonate preparation device, linked on-line to a Finnigan MAT 25 mass spectrometer. The analyses were performed at the
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Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research at Kiel University, Germany. Isotope values are presented in per mil (x) relative to the Pee Dee Belemnite (PDB) standard and NBS-20 standards were used between samples. The analytical precision was better than P 0.08x for N13 C and P 0.08x for N18 O. Isotope measurements are listed in Table 1. The equation from Bemis et al. (1998), T = 16.5^4.80 * (Nc 3Nw ), was used to calculate paleotemperature. Since most authors assume that extensive ice sheets did not exist during the Cretaceous, a N18 Ow of 31.22x estimated for preglacial seawater (Shackleton and Kennett, 1975) was used to estimate the paleotemperature in bottom water. Considering latitudinal di¡erences in surface water N18 Ow during the mid-Cretaceous (Poulsen et al., 1999a), a N18 Ow of 31.47x was used for planktonic foraminifera and bulk sediment. When calculating paleotemperatures, no vital e¡ects were considered.
3. Results 3.1. Biostratigraphy and carbon isotope stratigraphy We correlated the carbonate N13 C isotope curve at Site 551 to the N13 C isotope curve from the Eastbourne section in southern England (Fig. 4), since this is the most expanded and complete CTBI record in NW Europe (Paul et al., 1999). In the late Cenomanian, the Eastbourne carbonate N13 C values show ¢rst a slow increase, then a more rapid increase through the lower half of the Plenus Marls (Paul et al., 1999). The N13 C curve then shows two prominent peaks separated by a trough in the upper half of the Plenus Marls (Paul et al., 1999). The last occurrence of Rotalipora cushmani is within this interval. After the second
Fig. 3. Scanning electron micrographs of broken and unbroken tests of planktonic foraminifera from Site 551. Tests show excellent preservation and little recrystallization. Scale bar equals 100 Wm for specimens 1^5; and 30 Wm for specimen 6. (1) Rotalipora cushmani, Sample 80, 551-6R-2W, 138^141 cm. (2) Rotalipora greenhornensis, Sample 80, 551-6R-2W, 138^141 cm. (3) Dicarinella hagni, Sample 80, 551-5R-1W, 87^90 cm. (4) Dicarinella sp., Sample 80, 551-5R-1W, 87^90 cm. (5) Whiteinella aprica, 80, 551-6R-2W, 138^141 cm. (6) Whiteinella sp., Sample 80, 551-5R-1W, 87^90 cm.
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Fig. 4. Correlation of bulk sediment carbon isotope (N13 C) curves for Tarfaya Drill Hole S75 (Kuhnt et al., 2001), Eastbourne (Paul et al., 1999) and Site 551. Organic carbon was used for the carbon isotope measurements at Tarfaya Drill Hole S75, while carbonate was used for the carbon isotope measurements at the other two sites. The ¢rst occurrence of the nannofossil Quadrum gartneri is indicated in the Eastbourne curve. Rotaliporids are only present in Core 551-5R, their extinction level falls in the gap between Cores 551-5R and -6R.
peak, high carbon isotope values continue into the Holywell Marls, forming a plateau, and ¢nally carbon isotope values decrease gradually (Paul et al., 1999). In Eastbourne, the whole carbon isotope excursion is extended over about 17 m. The ammonite Watinoceras devonense Wright and Kennedy, which is used to de¢ne the base of the Turonian stage (Bengtson, 1996), is not found at Eastbourne. Therefore, the base of the Turonian stage is taken locally between the last occurrence of Neocardioceras juddii and the ¢rst appearance of Fagesia catinus (Gale, 1996). The nannofossil Quadrum gartneri ¢rst occurs at about 0.5 m above the last occurrence of N. juddii, close to the Cenomanian/Turonian boundary proposed by Gale (1996). With respect to the carbon isotope excursion, the Cenomanian/Turonian boundary is placed in the uppermost part of the N13 C plateau at Eastbourne. The presence of Rotalipora cushmani, R. deeckei
and R. greenhornensis and the absence of a carbon isotope excursion throughout Sections 551-6R-1 to -3 indicate a late Cenomanian age for this interval. The lower part of Core 551-5R, which includes the second carbon isotope buildup and peak, is consequently of latest Cenomanian age (Fig. 4). Sections 551-5R-1 to -2 can be correlated with the second carbon isotope peak and decline in the Plenus Marls of Eastbourne, based on the absence of Rotalipora and the presence of a carbon isotope peak and decline in this interval. 3.2. Benthic foraminiferal assemblages The carbonates below and above the black shales at Site 551 contain well preserved planktonic and benthic foraminiferal assemblages. The ratio of planktonic to benthic foraminifera remains extremely high, except within the black shale where both planktonic and benthic forami-
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Fig. 5. Stable isotope plots of bulk sediment, planktonic foraminifera and benthic foraminifera from Site 551. Temperatures are estimated from planktonic foraminiferal and bulk sediment N18 O values. N18 O temperature estimates are 1.2‡C higher for benthic foraminifera, because no correction for latitudinal di¡erences in Cretaceous sea water N18 Ow is applied.
nifera occur rarely. The well diversi¢ed benthic foraminiferal assemblages above and below the black shale commonly include Gavelinella dakotensis, Gyroidinoides lenticulus, Marssonella oxycona, Clavulinoides gaultinus, Praebulimina nannina, Tappanina laciniosa, Osangularia sp., Pleurostomella spp., Spiroplectinata spp. as well as various nodosariids. By contrast, the black shale usually contains only rare specimens of G. dakotensis, G. lenticulus, Osangularia sp. and T. laciniosa. 3.3. Oxygen isotopes in bulk sediment In Sections 551-6R-1 to -3 (138.67^142.30 me-
ter below sea £oor (mbsf)), the N18 O values show marked £uctuations around a mean value of 31.3x (Fig. 5). In Sections 551-5R-1 to -2 (132.56^134.66 mbsf), the N18 O values decrease from 32.2x at 135.66 mbsf to reach a minimum value of 32.9x at 134.76 mbsf, which corresponds to a N13 C peak. The N18 O values then generally increase above the black shale reaching 31.5xat 132.56 mbsf. Light and SEM microscope observations indicate that the bulk sediment generally consists of coccoliths (Fig. 6). To test whether the bulk signal was a¡ected by other components (such as fragments of foraminiferal tests), we analyzed parallel
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samples of ¢ne fractions 6 63 Wm. The di¡erence between bulk sediment and ¢ne fractions 6 63 Wm was found to be small. For a set of nine parallel samples, the mean di¡erence is 0.06x for N18 O and 0.14x for N13 C. Thus, it can be concluded that the isotope signal in the bulk sediment mainly represents the 6 63-Wm fraction (principally composed of coccoliths). 3.4. Oxygen isotopes in foraminifera In Core 551-6R, the planktonic foraminifera Hedbergella delrioensis show N18 O values varying between 31.6 and 32.05x with a mean value of 31.8x. Values for Rotalipora cushmani and R. greenhornensis are generally about 0.3x higher, and bulk sediment 0.3^0.8x higher than H. delrioensis (Fig. 5). In Core 551-5R, the N18 O values of Whiteinella aprica ¢rst decrease within the black shale, then show an overall increase, whereas the N18 O values for bulk sediment exhibit an overall decline. Above the black shale, Dicarinella hagni, W. aprica and Whiteinella archaeocretacea display N18 O values that are generally similar to the N18 O values in bulk sediment. In the upper part of the section, the W. aprica N18 O curve displays a minimum of 33.2x at 132.89 mbsf. The same pattern is evident in the N18 O curves of D. hagni, W. archaeocretacea and in the bulk sediment curve. In Core 551-6R, the benthic foraminifera Gyroidinoides lenticulus displays stable N18 O values around a mean of 30.7x (Fig. 5). In Core 551-5R, the Tappanina laciniosa record shows a small N18 O peak (31.6x) at 133.77 mbsf. 3.5. Carbon isotopes in bulk sediment In Sections 551-6R-1 to -3 (138.67^142.30 mbsf), the bulk sediment shows persistent N13 C values around a mean of 2.4x (Figs. 3 and 5). In Sections 551-5R-1 and -2 (132.56^135.66 mbsf), the N13 C values in bulk sediment increase from 3.3x at 135.66 mbsf through the black shale and peak to 3.9x at 134.76 mbsf. Above the black shale, the N13 C values decrease from 3.1x at 134.37 mbsf to 2.1x at 132.56 mbsf, except for a smaller peak (3.2x) at 132.89 mbsf.
Fig. 6. Scanning electron micrographs of bulk sediment fraction 6 63 Wm, showing rich coccolith content. The upper micrograph is from Sample 80-551-5R-1, 16^19 cm above the black shale, the middle micrograph is from Sample 80-5515R-cc, 4^7cm, in the uppermost part of the black shale and the lower micrograph is from Sample 80-551-6R-3, 60^63 cm below the black shale.
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As for oxygen isotopes, the di¡erence between bulk sediment and ¢ne fractions 6 63 Wm was found to be small (mean di¡erence of 0.14x for nine parallel measurements). 3.6. Carbon isotopes in foraminifera In Core 551-6R, the planktonic foraminifera Hedbergella delrioensis shows stable N13 C values around a mean value of 2.4x. Rotalipora cushmani, R. greenhornensis and bulk sediment samples show similar values. In Core 551-5R, the N13 C values for Whiteinella aprica, Dicarinella hagni and Whiteinella archaeocretacea are generally stable around a mean value of 2.6x. These planktonic foraminiferal N13 C values are approximately 1x lower than bulk sediment values within the black shale (Fig. 5). In Core 551-6R, the benthic foraminiferal N13 C values are more variable around a mean of 1.44x. The N13 C values in Tappanina laciniosa are relatively stable around 2x within the black shale. Above the black shale, the N13 C values in T. laciniosa decrease by about 0.5x. In the interval 551-5R-1, 6^9 cm to 551-5R-2, 48^50 cm (132.56^134.48 mbsf), the N13 C values in T. laciniosa are overall at least 0.5x lower than for planktonic foraminifera and bulk sediment. Within the black shale (134.76^135.63 cm), the di¡erence between benthic foraminiferal and bulk sediment N13 C values increases to 1.5x.
4. Discussion 4.1. The paleotemperature record of Site 551 and its importance for the reconstruction of the Cenomanian/Turonian Greenhouse climate 4.1.1. Sea surface temperature at Site 551 during the CTBI We reconstructed Cenomanian/Turonian Sea surface temperature (SST) values from three different sets of N18 O measurements: bulk sediment, Hedbergella delrioensis tests and a combination of Whiteinella aprica and Whiteinella archaeocretacea tests, as the latter two species appear to have had very similar habitat and isotope frac-
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tionation. Bulk sediment and 6 63-Wm fractions consist almost exclusively of coccoliths, thus their oxygen isotope values re£ect temperatures near the sea surface. However, the reliability of these data is limited by: (1) diagenetic alteration, (2) unknown vital e¡ects in oxygen isotope fractionation of Cretaceous coccolithophorids, and (3) extreme seasonality of coccolithophorids which commonly represent the population of a single plankton bloom with a duration of a few days. The planktonic foraminiferal genera Hedbergella and Whiteinella are commonly assumed to have lived predominantly at shallow depths in surface waters. This assumption is based on morphologic comparison with modern planktonic foraminiferal taxa (Hart, 1980), their distribution in marginal areas (Leckie, 1987) and stable isotope values (D’Hondt and Arthur, 1995; Huber et al., 1995). Whiteinella aprica exhibits N18 O values in Core 551-5R that are similar to those of bulk sediment. Thus, we regard paleotemperature estimates of 18.6^23.8‡C for W. aprica and 16.6^23.4‡C for bulk sediment as reasonably reliable for surface water. In contrast, Hedbergella delrioensis displays consistently lower N18 O values than coccolith bulk sediment in Core 551-6R. Thus, paleotemperature estimates for H. delrioensis (17.0^19.3‡C) are generally higher than for bulk sediment (13.8^ 21.4‡C). This o¡set in N18 O values between late Cenomanian surface dwelling planktonic foraminifera and coccoliths (bulk sediment) may be explained by the existence of a strong seasonal contrast in the late Cenomanian. Coccoliths may have been produced during the colder spring while planktonic foraminifera grew during the warmer summer. The similarity in N18 O values between planktonic foraminifera and bulk sediment in the latest Cenomanian^early Turonian (Core 551-5R) may indicate less pronounced seasonal temperature variations in surface water. Alternative explanations would be di¡erent vital e¡ects for W. aprica and H. delrioensis or signi¢cantly di¡erent composition of coccolith assemblages in Cores 551-5R and -6R. A clear warming trend from the latest Cenomanian to the earliest Turonian is observed. Average SST estimates, based on planktonic foraminiferal N18 O values corrected for latitudinal di¡erence in
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surface water N18 O (Poulsen et al., 1999a), rise from 17^19‡C in Core 551-6R (late Cenomanian) to a maximum of 23‡C in the latest Cenomanian and earliest Turonian, within and directly above the black shale. These results are in accordance with the study by Frakes (1999), which indicated a colder climate during the middle/late Cenomanian (Core 551-6R) and increasing temperatures with a temperature maximum in the latest Cenomanian and early Turonian (Core 551-5R). Using pristine planktonic foraminifera, Pearson et al. (2001) estimated warmer tropical surface water paleotemperatures for the Maastrichtian and mid-Eocene than had been estimated before. This ¢nding highlighted the importance of test preservation and the in£uence that secondary calcite had on paleotemperature estimates (Pearson et al., 2001). Even for our well preserved material from Site 551, a minor contribution from secondary calcite crystals cannot be excluded. Thus the absolute values for our temperature estimates should be treated with caution. 4.1.2. Comparison with records from other localities In the Eastbourne C/T section, situated at comparable paleolatitude as Site 551, most bulk sediment N18 O values are about 1x lower than at Site 551 (Paul et al., 1999). However, the sediments from Eastbourne experienced signi¢cant diagenesis and the N18 O values for chalk are approximately 1.0^1.5x more negative than those for adjacent marls, which were less in£uenced by diagenetic processes (Paul et al., 1999). When this diagenetic alteration is taken into account, the early Turonian SST estimates at Eastbourne are within the range of those estimated for Site 551. More reliable and less diagenetically altered is the Cenomanian/Turonian paleotemperature record based on oxygen isotopes of planktonic foraminifera from ODP Site 1050 at Blake Nose, at approximately 27‡N paleolatitude. However, the sedimentary record of the Cenomanian/Turonian transition is not complete at Site 1050. The second buildup and peak in the carbon isotope curve, documented at Eastbourne (Paul et al., 1999) and at Site 551 (Fig. 4), are not present in the Blake Nose record (Huber et al., 1999). Dur-
ing the recovery after the second carbon isotope buildup and plateau, planktonic foraminifera show similar N18 O and N13 C values at both Sites 1050 and 551. In the subtropical (15‡N paleolatitude) eastern North Atlantic, N18 O paleotemperatures between 30 and 35‡C were estimated in the Tarfaya coastal basin, southern Morocco for the interval representing the CTBI carbon isotope excursion (Luderer and Kuhnt, 1997). These temperatures may be overestimated due to diagenetic alteration, and are approximately 5‡C too warm when compared to Poulsen et al.’s (1999a) Parallel Ocean Climate Model simulations. However, well preserved Turonian planktonic foraminifera from the western tropical Atlantic also yielded estimated N18 O temperatures of at least 30^33‡C (Wilson et al., 2002). A late Cenomanian warming trend and a temperature maximum at approximately the extinction level of Rotalipora cushmani followed by stepwise cooling, which are evident in the Tarfaya record, are also found in the Blake Nose, Eastbourne and Goban Spur records. 4.1.3. Bottom-water temperatures Brass et al. (1982) proposed that evaporation in subtropical marginal seas and the subsequent sinking of warm and saline waters were important processes for the formation of deep water in the Cretaceous. Models for mid-Cretaceous ocean circulation indicate strong downward convection of warm and saline waters from subtropical areas in the eastern Tethys (Barron and Peterson, 1990). High density, warm and saline water may have also been generated in the Gulf of Mexico and contributed to intermediate and deep-water formation (Woo et al., 1992). Increases in bottomwater temperatures strongly in£uence the development of low oxygen conditions in the deep sea in two main ways: (1) it leads to increased metabolism in exothermic organisms, and (2) decreases the ability for water to dissolve oxygen. However, it is still a matter of vigorous discussion whether the presence of warm, anoxic, saline deep water alone without signi¢cantly increased surface productivity could cause deep-water anoxia and enhance the preservation of black shales in the deep sea.
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Deep-water paleotemperature estimates are more reliable since benthic foraminifera seem less prone to diagenetic alteration than planktonic foraminifera (Pearson et al., 2001). Thus, paleotemperature estimates of 12.9^16.9‡C (Gyroidinoides lenticulus) for Core 551-6R and 12.7^ 18.1‡C (Tappanina laciniosa) for Core 551-5R are considered as reasonably reliable for bottom water. Benthic foraminiferal N18 O values at Site 551 indicate no signi¢cant change in deep-water temperatures at the end of black shale deposition. Relatively uniform bottom-water temperatures through and immediately above the black shale (Fig. 5) suggest that elevated bottom-water temperatures were not the only reason for the development of low oxygen conditions at Site 551. It is likely that increased evaporation leading to the production of warm, highly saline and dense water during the CTBI occurred in Tethyan subtropical seas. On its way down, the warm and dense water plume may have entrained intermediate water and gradually become less saline and colder, before spreading out laterally as the Mediterranean water spreads into the Atlantic today. Thus, the warm climate during the CTBI may have stimulated the production of intermediate and deep water, in turn increasing vertical turnover and upwelling, and leading to increased productivity. If this scenario holds, the very warm climate at the CTBI could, at least partly, explain the increase in productivity and the deposition of organic rich sediments during this event. The warmest deep-water temperatures during the Cretaceous are recorded in the latest Cenomanian and early Turonian, when temperatures averaged 19‡C (Huber et al., 2002). At Site 551 (Core 6), Cenomanian deep-water paleotemperatures of approximately 13 to 16‡C are similar to or slightly lower than at Site 1050. Maximum temperatures of approximately 17‡C at Site 551 and 18‡C at Site 1050 follow the same trend during the isotope excursion. After the second carbon isotope peak bottom water, temperatures decreased from approximately 18 to 17‡C at Site 1050 (Huber et al., 1995) and from approximately 17 to 15‡C at Site 551. The slightly lower paleotemperatures at Site 551 are easily explained by the shallower pa-
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leodepth at Site 1050 (around 1200 m) than at Site 551 (around 2400 m). 4.2. Benthic foraminiferal distribution during the low oxygen event Within the black shale in Core 551-5R, the benthic foraminiferal assemblages display low diversity in contrast to intervals above and below the black shale. This is similar to the distribution in modern anoxic or close to anoxic marine sediments, where benthic foraminiferal assemblages generally show low diversity and low abundance (Bernhard et al., 1997; Gustafsson and Nordberg, 2000). Species within the black shale include Osangularia sp., Gyroidinoides lenticulus, Gavelinella dakotensis and Tappanina laciniosa, which are characteristic species of Albian and late Cenomanian black shales (Koutsoukos et al., 1990; Tronchetti and Grosheny, 1991; Kuhnt and Wiedmann, 1995; Erbacher et al., 1998; Holbourn et al., 2001). T. laciniosa and G. dakotensis, in particular, are characterisitc elements of late Cenomanian Tethyan and temperate shelf biofacies with enhanced carbon £ux and dysoxic bottom waters (Kuhnt and Wiedmann, 1995; Holbourn and Kuhnt, 2002). Site 551 represents the deepest occurrence of this typical shelf/slope high productivity assemblage (2400 m paleowater depth). This probably indicates an unusually high carbon £ux at this depth that only can be explained by considerably enhanced primary production. 4.3. Carbon £ux during the CTBI We observe a signi¢cant deviation between the N13 C curves of surface dwelling coccoliths (bulk sediment) and deeper dwelling non-keeled planktonic foraminifera within the CTBI black shale (Fig. 5). Above and below the black shale, planktonic foraminifera (Whiteinella aprica in the upper part of Core 551-5R and Hedbergella delrioensis in Core 551-6R) exhibit N13 C values similar to that of coccolith bulk sediment. However, within the black shale W. aprica and W. archaeocretacea have signi¢cantly lower N13 C values than bulk sediment, and show a minimum in N13 C values
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or no N13 C peak. A steep N13 C gradient through the upper water column during black shale deposition would best account for the deviation in N13 C signals, which can only be caused by signi¢cantly increased primary production and carbon £ux rates. Since both coccoliths and planktonic foraminifera live in the photic zone, the N13 C gradient must have extended from the water surface down to a few tenths of meters. The small depth extent of the N13 C gradient is also supported by the similarity in the oxygen isotope values of coccoliths and planktonic foraminifera. The similarity in the shape of the N13 C curves for coccoliths and benthic foraminifera suggests that the N13 C excursion was caused by global reservoir changes. However, the benthic N13 C excursion has a signi¢cantly lower amplitude (generally less than 1x) than the coccolith curve, which can be explained by the fact that the remineralization of large amounts of organic material signi¢cantly decreased N13 C values in sediment pore waters. Wilson and Norris (2001) proposed that a breakdown in water strati¢cation during the late Albian (OAE 1d) led to increased upwelling, increased production and subsequently to low oxygen bottom conditions and deposition of organic rich sediments. The situation at the CTBI may have been comparable. Weak water temperature strati¢cation would mean increased upmixing/upwelling of nutrient rich deepwater. This would lead to increased production and to the development of a steep N13 C surface water gradient, as well as increased export of organic material and subsequent development of anoxic or dysoxic conditions at the sea£oor during black shale deposition. Within the black shale at Site 551, we observe an unusual steep carbon isotope gradient between surface and bottom water, which can be accounted for by high productivity and an extremely active biological pump. High productivity is not possible in strongly strati¢ed water masses, since the uptransport of nutrients through the pycnocline is very limited. As Site 551 is far from the coastline, a signi¢cant nutrient supply from land is also not likely. At Sites 1050 and 551, planktonic and benthic foraminiferal N13 C values through the carbon isotope excursion are
comparable in magnitude. However, at Site 1050 planktonic foraminifera show a distinct N13 C peak through the black shale (Huber et al., 1999), which is not the case at Site 551 as mentioned earlier. It is therefore likely that the surface water N13 C gradient was not so steep at Site 1050, and primary production was possibly lower than at Site 551.
5. Conclusions The good preservation of calcareous microfossils together with the great paleodepth (about 2400 m) of DSDP Site 551 at Goban Spur (midlatitude NE Atlantic) provide a unique opportunity to use stable oxygen and carbon isotope measurements to reconstruct vertical temperature gradients and to assess the e⁄ciency of the biological pump in the water column during the CTBI. Stable isotope data of bulk sediment (coccolith) and well preserved tests of planktonic and benthic foraminifera allow to quantify surface and deepwater temperature changes that occurred during the CTBI. SSTs of about 18‡C and deep-water temperatures of about 14‡C indicate a low temperature gradient within the water column during the late Cenomanian. Following additional warming of at least 2‡C in both surface and deep water in the latest Cenomanian reaching a temperature maximum during black shale deposition, temperatures decreased again slightly in the early Turonian. Even in our well preserved material, a minor contribution of secondary calcite crystals cannot be excluded. Although absolute temperature values should be treated with some caution, temperature gradients are generally not in£uenced by this process. The CTBI carbon isotope excursion is prominent in the bulk (coccolith) signal, but less pronounced in the planktonic and benthic foraminiferal records. This di¡erence indicates a very steep 13 C gradient in the upper water column within the photic zone and a very e⁄cient biological pump during the CTBI carbon isotope excursion. We suggest signi¢cantly increased primary production in the uppermost water column and a shallow chlorophyll maximum within a water depth of
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few meters or tens of meters as the cause for this steep carbon isotope gradient. The temperature changes at Goban Spur exhibit the same trend of warming towards the CTBI black shale interval as at low latitude locations: ODP Site 1050 Blake Nose, and Tarfaya, Morocco. The latitudinal temperature gradient is in agreement with the (GCM) simulations by Poulsen et al. (1999a, b).
Acknowledgements We thank Dr. A. Niederbragt and Professor I. Premoli Silva for very helpful reviews. We also extend our thanks to Dr. Helmut Erlenkeuser at the Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research, to the sta¡ of the SEM Laboratory at the Institut fu«r Geowissenschaften, both at the Christian Albrechts University in Kiel and to the sta¡ of the ODP Core Repository in Bremen. This research used samples provided by the Ocean Drilling Program. The ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions Inc. We gratefully acknowledge ¢nancial support from the E.U. ¢nanced program ‘Rapid global change during the Cenomanian/Turonian OAE : Examination of a natural climatic experiment in earth history’, Grant HPRN-CT 199900055. The Deutsche Forschungsgemeinschaft is also gratefully acknowledged for ¢nancial support (Grants KU 649/7 and KU 649/8).
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