Ice age at the Middle–Late Jurassic transition?

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Earth and Planetary Science Letters 213 (2003) 205^220 www.elsevier.com/locate/epsl

Ice age at the Middle^Late Jurassic transition? G. Dromart a;
, J.-P. Garcia b , S. Picard c , F. Atrops a , C. Le¤cuyer a , S.M.F. Sheppard c a

c

Universite¤ de Lyon 1, UFR Sciences de la Terre, UMR CNRS 5125, 69622 Villeurbanne Cedex, France b Universite¤ de Bourgogne, UFR Sciences de la Terre, UMR CNRS 5561, 21000 Dijon, France Ecole Normale Supe¤rieure de Lyon, Laboratoire des Sciences de la Terre, UMR CNRS 5570, 69364 Lyon Cedex 07, France Received 20 September 2002; received in revised form 12 March 2003; accepted 5 May 2003

Abstract A detailed record of sea surface temperatures in the Northern Hemisphere based on migration of marine invertebrate fauna (ammonites) and isotopic thermometry (N18 O values of shark tooth enamel) indicates a severe cooling at the Middle^Late Jurassic transition (MLJT), about 160 Ma ago. The magnitude of refrigeration (1^3‡C for lower middle latitudes) and its coincidence in time with an abrupt global-scale fall of sea level documented through sequence stratigraphy are both suggestive of continental ice formation at this time. Ice sheets may have developed over the high-latitude mountainous regions of Far-East Russia. The drastic cooling just post-dated the Middle^Late Callovian widespread deposition of organic-rich marine sediments (e.g. northwestern Europe, Central Atlantic, and Arabian Peninsula). This thermal deterioration can thus be ascribed to a downdraw in atmospheric CO2 via enhanced organic carbon burial which acted as a negative feedback effect (i.e. the inverse greenhouse effect). The glacial episode of the MLJT climaxed in the Late Callovian, lasted about 2.6 Myr, and had a pronounced asymmetrical pattern composed of an abrupt (V0.8 Myr) temperature fall opposed to a long-term (V1.8 Myr), stepwise recovery. The glacial conditions at the MLJT reveal that atmospheric CO2 levels could have dropped temporarily to values lower than 500 ppmv during Mesozoic times. = 2003 Elsevier B.V. All rights reserved. Keywords: Jurassic; Paleoclimate; sea level; oxygen isotopes; ammonites

1. Introduction Jurassic climate has been considered for a long time as a typical ‘greenhouse climate’ with minimal equator-to-pole thermal gradients [1]. It assumes that such a general condition was related to high atmospheric concentrations of the green-

* Corresponding author.

house gas carbon dioxide. This view has been recently supported by estimates of the Jurassic [CO2 ] ranging between 1200 and 3000 ppmv, that is approximately 4^10 times modern (pre-industrial) values. The techniques for restoring ancient pCO2 include isotope records of pedogenic carbonate [2] and goethite [3], stomatal abundance from fossil leaves [4], and numerical simulations [5]. Here, we challenge the putative uniform and equable Jurassic climate with evidence for a severe, brief and global cooling that a¡ected

0012-821X / 03 / $ ^ see front matter = 2003 Elsevier B.V. All rights reserved. doi:10.1016/S0012-821X(03)00287-5

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Fig. 1. Migration towards subtropical latitudes of boreal ammonite fauna during the Late Callovian^Early Oxfordian epoch. Land contours modi¢ed after Smith et al. [7].

sea surface temperatures (SSTs) at the Middle^ Late Jurassic Transition (MLJT), around the Callovian^Oxfordian boundary, and was plausibly associated with the formation of subpolar continental ice. This view is derived from the convergence of data from paleobiogeography, O-isotopes (paleothermometry), and sequence stratigraphy (sea-level variations). Parameters of the climatic perturbation that are particularly noted are (a) the magnitude of SSTs changes, (b) comparison of the periods of initiation and termination, and (c) duration. In this paper, we also consider the mechanism that led to decline in temperature. Excess carbon burial in marine sediments results in a typical negative feedback e¡ect, i.e. the inverse greenhouse e¡ect.

2. Paleobiogeographic data The geographic distribution of ammonites is a qualitative proxy for the SST evolution because most of these marine nekton are considered by Martill et al. [6] to have inhabited the upper parts of the water column. During most of the Mesozoic, in the Northern Hemisphere, ammonites

were separated into two distinct realms: the cool, boreal and the warm, Tethyan domains. Boreal ammonites had invaded low-latitude seas by the end of Middle Jurassic times (Fig. 1) and the latitudinal magnitude of the migration varies from 15‡ (NW Paci¢c) to 30‡ (NE Paci¢c) (Table 1). In western Europe, the spreading pattern of these marine invertebrate faunas can be reconstructed in detail (Fig. 2). Throughout the Bathonian, boreal ammonites (i.e. Cadoceras spp. ) were restricted to northern Europe (Shetlands and West Greenland) [12]. From latest Bathonian to Middle Callovian times, boreal fauna show brief and limited southward incursions [13]. Then boreal Kosmoceratidae steadily settled in southern Europe during Late Callovian times, reaching south Portugal (Algarve) during the lamberti Zone [14]. A subsequent wave of Boreal ammonites (i.e. Cardioceratidae) penetrated SE France during the latest Callovian (lamberti Subzone) [15], and reached northwestern Africa during earliest Oxfordian times [10]. Con¢rmation that marine boreal fauna of northwestern Europe migrated southwards in response to a decline of SSTs comes from paleo£oral data which indicate a general climatic modi¢-

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cation in the area. Change of palynomorph associations observed in the Upper Callovian of the North Sea [17] and the occurrence of Xenoxylon plants in the Upper Callovian and Lower Oxfordian of Germany and France [18] are both diagnostic of the onset of cool and humid conditions on the N Europe hinterland. The remaining issue is to determine the magnitude of the SST decline at mid-latitudes.

3. Isotopic data Prior to the Early Cretaceous radiation of planktonic foraminifers, ectothermal vertebrates living in surface waters can be used as paleorecorders of SSTs. The N18 O value of biogenic apatite, specially enamel whose high degree of mineralization makes it more resistant to post-depositional isotopic exchange, varies with both the N18 O value and temperature of water. Empirical fractionation equations have been established from modern ¢sh, yielding a single temperature-dependent curve for all ¢sh species [19^21]. This intrinsic property of the phosphate^oxygen^water system can thus be applied to extinct species to infer ancient temperatures, assuming a certain N18 O value for the ambient water. The N18 O composition of chondrichtyan ¢sh teeth (i.e. Asteracanthus spp. also referred to as Strophodus spp. ) coming from the continental shelf of the northwestern Tethys (eastern France and Switzerland) have been analyzed (Table 2), provided that this shark was a surface or near-surface dweller. The large hybodont shark Asteracanthus spp. has a durophagous dentition able to crush any

mollusk shell. It may have been a predator of surface-living ammonites and benthic bivalves. The teeth of Asteracanthus have been found in a variety of sedimentary facies in Jurassic platforms of western Europe, ranging from very shallow, restricted facies to outer, o¡shore facies (above the storm-weather wave-base). This is suggestive of a ¢sh tolerating a variety of environments whose depth did not exceed 100 m. To characterize the living depth of Jurassic ¢sh, Picard et al. [23] have compared N18 O values of ¢sh tooth enamel to a data set on co-existing calcitic brachiopods which are typical elements of the benthos. Results reveal temperatures ranging from 20 to 27‡C for subtropical (V30^34‡N of latitude) surface seawater in the late Middle Jurassic (Fig. 2), assuming for the moment a N18 O seawater of 31x (SMOW), that is for an ice-free world reference. These isotopic temperatures are in agreement with both the assumption that the sharks inhabited super¢cial waters and the host rocks contain corals and oolites, typical of ‘tropical carbonates’, the development of which in modern seas requires a mean annual temperature of at least 20‡C [24]. The Late Callovian epoch in the area is marked by an apparent drop of V5‡C of surface waters, cooling down to V17^19‡C. It should be noted that: (1) the belemnite N18 O data from the MLJT of the United Kingdom as well as Russia show a similar abrupt shift to heavier values, interpreted as due to climatic deterioration [25]; 2) the seawater cooling indicated by the O-isotope paleothermometry coincided exactly with the massive southward movement of boreal ammonites (Fig. 2). In the very late Middle Jurassic, the change

Table 1 Distribution of boreal ammonites over the Northern Hemisphere at the MLJT Province Species Maximum South Location Paleolatitudea Arrival Time

NW Paci¢c not speci¢ed

Latitudinal Migration Span Reference

15‡C Sey et al. [8]

a

Vladivostok 60‡N Callovian-Oxfordian

207

NE Paci¢c W Eurafrica Cardioceras cordiform Cardioceras £exuosum Northern Utah W Morocco Agadir Area 30‡N 25‡N Early Oxfordian Early Oxfordian mariae Zone 30‡C 30‡C Poulton et al. [9] Ambroggi [10]

Estimated from Smith et al. [7].

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Eurasia Quenstedtoceras lamberti Northern Iran Central Alborz 30‡C Late Callovian lamberti Zone 20‡C Seyed-Emani et al. [11]

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Fig. 2. Covariation across the MLJT in western Europe of paleogeographic expansion of ammonites and SSTs. Distribution of ammonites is derived from our own observations [F.A., G.D.] plus references [10,12^16]. Paleolatitudes are derived from Smith et al. [7]. SSTs are derived from new O-isotope data of apatite enamel from shark teeth (Table 2). Light-colored triangles for the Upper Callovian and Lower Oxfordian are derived by assuming a N18 Osw value of 30.5 to 0x during glacial maximum, giving glacial isotopic temperatures 2^4‡C higher than the model values with N18 Osw = 31 (triangles joined by arrows correspond to the same tooth and same N18 O value, but di¡erent N18 Osw ).

recorded for subtropical SSTs is suggestive of the formation of high latitude continental ice because the magnitude of the temperature decline (5‡C) is similar to estimates of the temperature shift for the subtropical zones between interglacial to glacial Quaternary episodes [26]. Additional indication for polar cooling throughout the Early Oxfordian may come from the steep latitudinal gradient for SSTs (0.6‡C per degree of latitude) tentatively calculated between northern Eurasia and western Europe. For the Lower Oxfordian, mean seawater temperatures calculated from N18 O values of belemnites (carbonate system

[27]) of the Moscow area (47‡N) and ¢sh (phosphate system, Table 2) of East France (35‡N) are 7.3 and 18.5‡C, respectively. A simple deviation of 11.2‡C can be proposed assuming no latitudinal variation in the N18 O value of seawaters. Moreover, Anderson et al. [28] have noted that the belemnites of the Peterborough Member (Pt Mb; Middle Callovian of southern England) yield paleotemperatures similar to those of benthic bivalves, and 3^4‡C cooler than paleotemperatures derived from ammonites inhabiting the super¢cial waters. Application of such a living-depth correction yields a real di¡erence of 7.5‡C for SSTs

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Table 2 Oxygen-isotope composition of apatite enamel from shark teeth of the Middle and Upper Jurassic in western Europe (East France, Switzerland)a Taxa

Locality

Stratigraphy

Standard zonation

Marker-bed

Age

N18 O (PO4 )b

Tc

(ammonite zones)

(brachiopod)

(Ma)

(x) (SMOW)

(‡C) (N18 Osw = 31)

(ammonite subzones)

L. Kimmeridgian U. Oxfordian

[PlanulaPlatynota] Planula

154.0

19.7

22.4

154.2

19.7d

22.4

Jura, Salins Jura, Ornans

L. Oxfordian L. Oxfordian

158.3 158.8

20.4 20.9

19.4 17.2

Asteracanthus

Arvel (CH)

159.4

20.5

19.0

Asteracanthus Asteracanthus Asteracanthus Asteracanthus tenuis Strophodus reticulatus Strophodus Asteracanthus Asteracanthus

Co“te d’Or, Talant Co“te d’Or, Prusly Co“te d’Or, Etrochey Co“te d’Or, Authoison Co“te d’Or, Etrochey Jura, Ornans Yonne, Brion (well) Co“te d’Or, Colline St-Anne Co“te d’Or, Ladoix Co“te d’Or, Merey

Callovian/ Oxfordian U. Callovian U. Callovian M. Callovian M. Callovian L. Callovian L. Callovian L. Callovian L. Callovian

Cordatum Mariae/ Cordatum Lamberti/Mariae Athleta Athleta Coronatum Jason Calloviense Koenigi Herveyi Herveyi

Torqui Kalli Divio Divio

160.2 160.5 161.1 161.5 162.2 163.3 163.9 163.9

19.5d 18.9 19.2 19.1 18.8d 19.6 20.2 19.8

23.1 25.7 24.6 25.2 26.3 22.7 20.4 22

L. Callovian BathonianCallovian U. Bathonian

Herveyi DiscusHerveyi Discus

Terebratus DiscusKlepperi Discus

Divio [Eudesia-Divio]

163.9 164.2

19.0 19.1

25.4 25.0

Eudesia

164.6

19.1d

25.0

Discus Eudesia Hannoveranus Digo

164.6 165.3

20.7 20.2

19.9 20.1

166.8

19.6

22.8

167.6

19.7d

22.4

Asteracanthus orna-tissimus Asteracanthus Asteracanthus EPSL 6698 5-8-03 Cyaan Magenta Geel Zwart

Asteracanthus Strophodus reticulatus

U. Bathonian U. Bathonian

Discus Orbis

Asteracanthus

Co“te d’Or, Les Perrie'res (Dijon) Yonne, Jaulges (well) Co“te d’Or, NuitsSt-Georges Ain, Nantua

M. Bathonian

Morisi

Strophodus

Co“te d’Or, Vanvey

M. Bathonian

Progracilis

Asteracanthus Asteracanthus Asteracanthus

Phaeinum Grossouvrei Enodatum Terebratus Terebratus

Oxoniensis

[MultiplicataConcinna] [GlobataMultiplicata]

G. Dromart et al. / Earth and Planetary Science Letters 213 (2003) 205^220

Sao“ne and Loire, Flace¤-les-Ma“con Yonne, Tonnerre

Asteracanthus

a All samples come from open marine environments, i.e. shoreface and o¡shore. Samples coming from protected environments have not been included to prevent any mistake in the temperature conversion due to the modi¢cation of the N18 Osw by excessive evaporation. b O-isotope analysis at Ecole normale supe¤rieure de Lyon. c Temperatures of seawater calculated according to the procedure described by Longinelli and Nuti [19] and Le¤cuyer et al. [21], assuming N18 O seawater = 31x (SMOW). d Also analyzed for N18 O (CO3 ) in apatite [22]. All samples are consistent with isotopic equilibrium between phosphate and carbonate.

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EPSL 6698 5-8-03 Cyaan Magenta Geel Zwart Fig. 3. Intercontinental extension and expression of a regressive event at the MLJT. The sea-level fall at the end of Callovian times is expressed by a variety of sedimentological evidence depending on depositional settings: basin fans in deep waters; by-pass (hiati) and ravinement (erosion) on shelves; valley incision and continental deposits (soils, karst, eolian and £uvial systems) sandwiched between marine sediments or near-shore environments. Biostratigraphic allocations of sections are all based on ammonite assemblages but have di¡erent resolution (subzone to substage levels). Downward extension of the stratigraphic hiatus into the Callovian is variable because of more or less pronounced erosion of former deposited strata associated with the sea-level fall. Data from [8,9,30,31,38^54].

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between Russia and France. A gradient of V0.6‡C per degree of latitude can thus be inferred. Even if this is tentative, because of the limited number of ¢sh samples and the scatter of the belemnite N18 O data [27], note that the estimated gradient for the Jurassic at least equals the modern gradient range, that is 0.2 to 0.5‡C per degree of latitude for the equivalent latitudinal belt of the Northern Hemisphere.

4. Sequence stratigraphy The formation of continental ice is the most e⁄cient mechanism to account for rapid, globalscale sea-level change. The publication of the chronology of £uctuating sea levels since the Triassic by Haq et al. [29] has prompted a large amount of stratigraphic documentation to check and re¢ne the proposed curve. Several versions of sea-level curve have then been proposed for the Middle^Upper Jurassic, based on various regional stratigraphic data [30^36]. While good agreement exists between these curves for some intervals, signi¢cant discrepancy appears for others. Notably, no consensus exists on the eustatic history at the MLJT, certainly re£ecting, on a regional scale, variations in the tectonic and sedimentary runo¡ contributions to the relative sea level. Here, we review sea-level variation at the MLJT on the basis of conventional facies and stratal analysis of depositional sequences. Because intercontinental correlation is required to detect eustatic phenomena, considered sections come from diverse worldwide localities, for which detailed and reliable biostratigraphic information is available, have been considered. The data base derives from our own observations (i.e. France, Algeria, Oman, Canada) and published stratigraphic columns (Fig. 3). After a well-established worldwide transgressive event that culminated during the Middle Callovian [37], there was a late Middle Jurassic sealevel lowstand, suggested by a variety of sedimentological features (Fig. 3): deposition of deep-sea fans (calcareous turbidites in Central Atlantic and Middle East (Oman); siliciclastic turbidites in the North Sea and North Algeria; seaward migration

211

of shorelines (progradation of shoreface sandstones all around the North Sea); valley incision in South Portugal and NE Saudi Arabia; subaerial exposures of marine sediments (paleokarsts in Portugal, Israel and northern Oman) ; alluvial fans in Alaska and Argentina). While regional sea-level fall is readily detected by analysis of sedimentary successions, dating and correlation over wide regions requires the presence of appropriate fossils. Jurassic £oodings (sea-level rises) can be dated and correlated based on ammonite assemblages. In contrast, regressive counterparts bear larger age uncertainty because they are associated with erosion and subaerial exposition, and with littoral and £uvial deposits which have no appropriate fossil record for dating. Apparently, there is no dating discrepancy between the given sections (Fig. 3). The uppermost Callovian marine level preserved beneath the MLJT disconformity is the henrici Subzone in age (Iberian ranges). Conversely, the ¢rst transgressive deposits above the MLJT disconformity is the lamberti Zone or lamberti Subzone in age (e.g. Paris Basin). From this data set, it is inferred that the sea-level minimum was reached at the henrici and lamberti Subzone transition in the latest Callovian. The sea-level fall at the MLJT can be viewed as globally well documented, of eustatic origin, with a maximum in the Late Callovian. The implication of this is that ice cap formation and maximum decline of SSTs should have occurred during the Late Callovian and not the Early Oxfordian, as suggested by the optimal southern extent of boreal ammonites in the northwestern Tethys (Table 1). This con£ict may only be apparent in that inappropriate depositional conditions (high sedimentation rate or erosional hiati) may have precluded the ¢nding of boreal ammonites in the Upper Callovian of southern areas. In addition, it should be noted that the boreal Cardioceratidae was very proli¢c in SE France during the latest Callovian and dropped in relative abundance across the Callovian^Oxfordian boundary (50% of ammonites down to 5^ 15%) [15]. The magnitude of sea-level fall can be estimated at several tens of meters (40^80 m) from depositional facies contrast in Portugal [39]. In the west-

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ern Lusitanian basin, Middle Jurassic outer ramp facies composed of interbedded marls and limestones with ammonites, brachiopods and bivalves are directly overlain by Middle Oxfordian freshwater/brackish facies including lignitic clays and marls with non-marine ostracods, charophytes, desiccation cracks, and pedogentic carbonate.

5. High-latitude continental ice The abrupt and global Late Callovian sea-level fall coincides remarkably with the decline of SSTs (Fig. 2). Such a close covariation is diagnostic of the growth of continental ice at that time. The remaining issue is ‘where was the ice hiding?’. Necessarily, the ice was stored somewhere on continents because £oating ice shelves do not change sea level. Jurassic global paleogeographic reconstructions show a number of landmasses rather than a major landmass centered about one or both polar positions [7], so that ice was probably not concentrated in a large, single mass. There are a number of supporting observations that ice may have formed around the northern pole: (1) high-latitude freezing conditions apparently have left direct and typical imprint in the form of glendonite (i.e. star-shaped carbonate pseudomorphs after the polymorph of CaCO3 , ikaite which crystallizes in subaqueous conditions at temperatures close to 0‡C [55]) reported from the Callovian-Oxfordian sediments of northeastern Asia [56]; (2) Middle and Late Jurassic plate con¢gurations of the Arctic reveal that continental blocks were drifting just around the pole [57], some of them being aerially exposed at the MLJT (Chukotka, Verkhoyanye, Olomon provinces) [8]; and (3) geological data also indicate that the (Bathonian^Oxfordian) interval in the far east of Siberia (Yablonovoi^Stanovoj Mountains) was marked by crustal folding and uplifting [8], probably leading to high-latitude mountainous regions. Finally, it should be noted that whatever the details of the con¢guration of the high-latitude continental province, important landmass surfaces were restricted inside the northern paleopolar circle, in which the modern permafrost is continuous, has high ice content, and is com-

monly more than 500 m thick [58]. Ice could have been present over the southern paleopolar regions during parts of the Jurassic [59]. Thus, it is quite plausible that continental ice may have been locked up in northern and southern highlatitude ice sheets and permafrost. Conversely, the formation of 16 O-rich ice sheets has implication for O-isotope paleothermometry because it makes the N18 O seawater value rise. The amount of ice can be derived from the magnitude of sea-level fall at the MLJT estimated at several tens of meters (40^80 m). On the basis of simple linear interpolation between the Last Glacial Maximum and present ice volumes [60], a 40^80 m fall would require an ice accumulation of 6^12 Mkm3 . Surfaces of the Chukotka and Verkhoyanye blocks, and Yablonovoi^Stanovoj Mountain Ranges are estimated at 1 Mkm2 each. The growth of 2^3-km-thick ice sheets over these areas (because of subpolar position or high elevation in high latitude) would result in the formation of a total ice volume of 6^9 Mkm3 . Subaerially exposed surfaces of Central Siberia above the paleo-Arctic circle were about 3 Mkm2 , permitting only an ice volume of 0.3 Mkm3 to be stored in the subsurface if we admit an ice content of 20% for a 500-m-thick permafrost. Because both the volume and surface area of the Late Callovian ice caps are estimated to be small relative to the Present ice caps, a N18 O of 330x is taken rather than the current model value of 350x [61]. This implies a N18 Osw value of 30.5 to 0x during glacial maximum, giving glacial isotopic temperatures 2^4‡C higher than the model values with N18 Osw = 31. Accordingly, the net cooling of 1^3‡C of amplitude recorded for the subtropical surface seawaters of the MLJT is still comparable to estimates for Quaternary interglacial to glacial episodes.

6. Inverse greenhouse e¡ect The ¢nal question is related to the mechanism that governed the Earth’s surface refrigeration over the MLJT. On the million-year time scale, there are two possible processes responsible for the uptake of CO2 from the atmosphere : (1)

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weathering of silicate rocks; (2) production and burial of organic matter. The atmospheric CO2 is converted during chemical weathering of silicate and carbonate rocks into dissolved HCO3 3 which is transferred to oceans by rivers and precipitated there as carbonate minerals. This mechanism can a priori be ruled out since there is no evidence of correlative enhanced carbonate deposition. On the contrary, the Middle Callovian is marked by a withdrawal of carbonate platforms (i.e. drowning and constriction) from the subtropical latitudes of the northern Hemisphere, in western Europe (Paris Basin) [62^63] and eastern North America (o¡shore Nova Scotia) [64]. Conversely, Middle Callovian marine organicrich sequences (i.e. bituminous shales) have been

213

reported from many low- and mid-latitude places all around the world (Table 3). The reference sections for the Callovian organic-rich deposits are the Pt Mb of the Oxford Clay Formation in England [65], subunit 7c of the Unnamed Formation of DSDP Site 534A in Central Atlantic [66], and the Tuwaiq Mountain Formation in Saudi Arabia [69]. The Pt Mb, 17 m thick at Peterborough, typically comprises dark greenish gray, shelly mudstone, with a minor cyclicity involving the alternation of dark gray, very ¢ssile units and ¢rmer, blocky, slightly paler mudstones. The Pt Mb has been allocated to the uppermost Lower Callovian ^ lowermost Upper Callovian range on the basis of ammonite biostratigraphy. The total organic carbon (TOC) content of the Pt Mb is up to 16.6% and mostly above 3 wt% (Fig. 4). Ma-

Table 3 Geographic distribution and composition of Callovian, organic-rich marine sediments Province

Location

Lithostratigraphy

Lithology (TOC max)

South England

Peterborough

Peterborough Mb

¢ssile, dark gray mudstone (16.6%)

Central Atlantic

ODP 534 A

East Paris Basin

HTM 102 well

West Scotland

Skye Island

Unamed Fm, 7c to 7e units Argiles de la Woe«vre Fm Dunans Shales Mb

West Greece

Ionan zone

Kurdistan

Urmia lake

Saudi Arabia

Khurais

Tuwaiq Mountain Fm

laminated mud/ packstone (13%)

Philippines

Mindoro

Mansalay Fm

Central Oregon

Izee area

South Mexico

Guerrero, Cualac DSDP 511

black shales, carbonaceous sands organic-rich M.-L. Callovian mudstone bituminous shales Callovian laminated shales

East Falkland

Epirus Radiolarites

Age

late E. to early L. Callovian (enodatum to phaenium SZ) laminated claystone M. to L. Callovian (3.3%) green-colored M. Callovien claystone (1%) (jason Z) laminated E.-M. Callovian bituminous shales (10.4%) radiolarites (8.6%) Callovian bituminous shales

M. Callovian (anceps Z) late M. to early L. Callovian (coronatum to athleta Z) E.-M. Callovian

M. Callovian (coronatum Z)

TOC: total organic carbon (wt%). Estimated from Smith et al. [7].

a

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Paleolatitudea

Reference

40‡N

Kenig et al. [65]

15‡N

Sheridan and Gradstein [66] Landais and Elie [67] Morton [44]

35‡N 45‡N

20‡N 30‡N

Danelian and Baudin [68] Arkell [45]

5‡S

Carrigan et al. [69]

30‡N

Andal [47]

30‡N

Poulton [9]

10‡N

Arkell [45]

60‡S

Ludwig and Krasheninnikov [70]

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rine algal matter of phytoplanktonic origin dominates the organic-rich shales. In the Central Atlantic, the Middle Callovian sequence is 8 m thick, and is composed of a laminated, greenish black to olive black claystone. It has been assigned to the Middle and Upper Callovian from the dino£agellate and calcareous nannoplankton assemblages. The TOC values are 1.1^3.3 wt% (Table 4). The organic matter is terrestrial with

a marine contribution when the TOC is high. In the southern Arabian Basin, the organic-rich interval of the Tuwaiq Formation is up to 150 m thick. The organic matter is dominated by lamalginite with high hydrogen indices. The lamalginite consists of leiosheres which form ¢ne, but very distinctive laminae in peloidal carbonate packstone. The TOC content averages 3 wt% with values up to 13 wt%. The Tuwaiq Formation has

Table 4 Pyrolysis data for Callovian and Oxfordian sediments of Central Atlantic (DSDP 534A) Subunit 534A

Position (core/section/ cm)

Sub-bottom depth (m)

Stratigraphy

Oxfordian Oxfordian Oxfordian Oxfordian Oxfordian (base) Oxfordian (base) Callovian/Oxfordian Callovian/Oxfordian Upper Callovian Upper Callovian Upper Callovian Upper Callovian Middle/Upper Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian Middle Callovian

7a 7a 7a 7a 7b 7b 7c 7c 7c 7c 7c 7c 7c

112/1/106a 115/1/34 115/1/36 115/1/38 119/1/111 119/1/120 122/1/14c 122/1/35a 122/2/46 123/2/79c 123/2/99c 123/3/20a 124/1/59a

1505.56 1531.84 1531.86 1531.88 1579.61 1579.70 1590.14 1590.35 1591.96 1592.29 1592.49 1597.20 1604.09

7c 7c 7c 7d 7d 7d 7d 7d 7d 7d 7e 7e 7e 7e 7e 7e 7e 7e

125/2/29 125/2/100a 125/3/100c 125/4/70c 125/5/44 125/6/40 126/2/30c 126/2/76a 126/2/118 126/3/16 126/4/20 127/1/05 127/1/120 127/1/130 127/2/36 127/2/37 127/2/42 127/3/74

1614.29 1615.00 1616.5 1617.70 1618.94 1620.40 1623.36 1623.76 1624.18 1624.66 1626.20 1630.55 1631.70 1631.80 1632.36 1632.37 1632.47 1634.24

Sediment

TOC

Tmax

IH

(%)

(‡C)

(mg HC/g TOC)

claystone dark reddish claystone olive greenish claystone dusky reddish claystone olive greenish claystone black claystone marly limestone (turbiditic) marly limestone (turbiditic) dark greenish, silty claystone marly limestone (turbiditic) dark claystone calcareous silty claystone marly limestone (turbiditic)

0.2b 0.03 0.02 0.02 0.15 0.03 0.38 0.5b 0.49 0.47 1.30 0.9b 0.4b

dark greenish, silty claystone dark greenish claystone black claystone black claystone dark greenish claystone black greenish claystone calcaceous claystone claystone dark greenish claystone dark greenish claystone calcareous claystone dark brownish claystone dark greenish claystone black claystone dark greenish claystone black, mm-thick lamina dark brownish claystone dark greenish claystone

1.97 1.8b 1.8 2.40 3.31 3.28 1.10 3.09b 2.68 1.44 0.16 0.11 1.06 1.29 0.4 3.11 0.05 0.35

0

423

46

429 429

70 106

429 433

102 83

435 431

186 161

433 434

226 237

430 435 434

221 205 60

420 432 434 428 432 434 426

63 117 121 100 270 20 94

Stratigraphy based on dino£agellates [75]. Rock-Eval pyrolysis data obtained at Institut Franc[ais de Pe¤trole. TOC: total organic carbon. Tmax : Pyrolysis thermal maximum, i.e. thermally immature organic matter (Tmax 6 435‡C). IH: pyrolysis hydrogen index, i.e. mixture of continental (IH 6 100) and marine (IH s 100) organic matter. a Data from Herbin et al. [74]. b LECO analysis. c Data from Sheridan and Gradstein [66].

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215

Fig. 4. Covariation of the organic matter content of Callovian^Oxfordian marine sediments of western Tethys and SSTs of northeastern and western Europe. The TOC is derived from pyrolysis analysis [71]. This data base is a compilation of new data (Central Atlantic, ODP 534A, Table 4) and previously published values [65,67,72]. Evolution of seawater temperatures in NE Europe (W Russian platform) calculated through Epstein’s equation [73] from O-isotopic composition of belemnite rostra [27]. Light gray values in the Upper Callovian and Lower Oxfordian (joined by arrows) are derived by assuming a N18 Osw value of 30.5 to 0x during a glacial maximum, giving glacial isotopic temperatures 2^4‡C higher than the model values with N18 Osw = 31. The henrici/lamberti Subzone transition corresponds to SST minimum as inferred from sea-level lowstand.

been assigned to the Middle and Upper Callovian in Arabia on the basis of ammonite biostratigraphy [52,76]. Deposition of organic-rich layers occurred in both epeiric seas (western Europe and Arabian Peninsula) and narrow deep troughs (Atlantic seaway and Mediterranean area) throughout the Middle Callovian ^ early Late Callovian interval. Supporting evidence for a global enhanced organic carbon burial comes from the carbon-isotope data from macrofossils from the Oxford Clay. Relatively high values of N13 C (5^ 5.5x) are recorded from bivalves and ammonites in the Middle Callovian and lowermost Upper Callovian (phaeinum Subzone) [28]. O-isotope records clearly show that the deposition of Middle Callovian organic-rich shales coincided in time with the thermal optimum of seawater (Fig. 4). We do not know however why

carbon burial was transiently increased. Possibilities are: (a) enhanced preservation of organic matter due to diminished O2 solubility in warmer water, and/or (b) excessive organic productivity supported by high [CO2ðatmÞ ]. The fact that the organic-rich sediments were preferentially deposited in medium- to low-latitude areas (Table 3) provides support for the ¢rst hypothesis. The seawater cooling that began in the early Late Callovian, athleta Zone, set in immediately following the episode of widespread deposition of organic carbon (Fig. 4). There is a time-gap of about 0.8 Myr [77] between the termination of the organic rich deposition in central England (phaeinum/proniae Subzone boundary) and sealevel minimum (henrici/lamberti Subzone boundary). This lag-time, which corresponds to the time required to build continental ice sheets,

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G. Dromart et al. / Earth and Planetary Science Letters 213 (2003) 205^220

may have been much shorter (i.e. 0.2^0.3 Myr) if we consider that organic-rich sedimentation ceased at the end of the athleta Zone in Arabia [52]. This fairly abrupt thermal deterioration can thus be ascribed to a downdraw in atmospheric CO2 via enhanced organic carbon sequestration which acted as a negative feedback e¡ect, i.e. the inverse greenhouse e¡ect. Such an impact of the Corg storage on climate supports the general assumption that on geologic time scales, the major control on global climate was the abundance of CO2 in the atmosphere.

7. Ice waning The sea level resumed rising in the latest Callovian and the ensuing positive trend was marked by a number of transgressive blips during the Early Oxfordian, i.e. scarburgense and bukowski Subzones, and the early Middle Oxfordian, i.e. the antecedens Subzone (Fig. 3). The sea-level behavior at the MLJT yields a typical ‘saw-tooth’ pattern, with a sharp fall of sea level followed by a long-term, stepwise rise. This asymmetrical pattern suggests that ice froze up as a single and geologically rapid process whereas it melted down at variable rates. The ammonite distribution (Fig. 2), paleo£oral assemblages [17], and belemnite N18 O data [25] from the Upper Jurassic of western Europe together document that temperatures eventually increased during the early Middle Oxfordian. The sudden retreat of Cardioceratidae to the boreal province with reciprocal massive invasion of Tethyan forms, i. e. Oppeliidae (Taramelliceras), marked in southeastern France the passage between the vertebrale and antecedens Subzones of the early Middle Oxfordian [16]. Consistently, this event coincided in time with the settlement of oolitic/coral carbonate deposits in northern areas : northwestern Switzerland (Saint Ursanne Fm [78]) and southern England (Osmington oolite and Coral-rag [79]). The sharp change of ammonite hegemony suggests an abrupt warming of SSTs during the Middle Oxfordian in western Europe. Unfortunately, vertebrate (i.e. Asteracanthus) O-isotope data are too sparse for checking

the timing and estimating the magnitude of this thermal recovery (Fig. 2). It is thus inferred that the glacial episode of the MLJT lasted 2.6 Myr, from the relation of the biostratigraphic data to the geochronologic scale [77], and had a pronounced asymmetrical pattern composed of an abrupt (0.8 Myr at the utmost) temperature fall opposed to a long-term (1.8 Myr), stepwise recovery. However, the Earth’s climate as a whole did not return to its former status. The MLJT is one of the major turning points of the climatic history as suggested by the major coeval modi¢cation in the distribution of the zones of vegetation [80]. The Late Jurassic climate was markedly di¡erent in that arid conditions prevailed all over the (Eur)Asian hinterland [80] whereas chilly and humid conditions persisted after the MLJT in the North Paci¢c region [81], leading to strong contrasts in the longitudinal thermal conditions in the Northern Hemisphere.

8. Conclusion The time di¡erence between the proposed maxima of cooling of seawater in the Early Oxfordian and drop in sea level in the Late Callovian may be more apparent than real. Because of the relatively limited number of samples available from the Upper Callovian, and hence few O-isotope analyses, the SST minimum cannot be pinpointed with the necessary very high precision. Also, satisfactory de¢nition of its magnitude is not helped by the currently inevitable poor control on N18 O seawater. Nevertheless, evidence for the subpolar glaciation, the proposed mechanism responsible for the global change in oceanographic conditions (SSTs, sea level, N18 O) and redistribution of fauna at the MLJT, is consistent with presently available geochemical and geological information. In conclusion, we suggest that in contrast to CH4 outbursts from gas hydrate dissociation [82] and to CO2 release by continental £ood basalts [83] which may have transiently raised the pCO2 during the Jurassic, enhanced sequestration of organic carbon in sediments may have caused ‘icehouse’ interludes, plausibly associated with abrupt and brief falls of the CO2 levels to values

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lower than 500 ppmv. On the scale of 1 Myr, the self-regulation of CO2 level and surface temperature via the production/preservation of organic matter appears to be a signi¢cant driving mechanism of global climate change. The documentation of this ‘cold snap’ substantiates the designation by Frakes et al. [84] of the Middle Jurassic^ Early Cretaceous interval as a Cool Mode of the Phanerozoic. Finally, this ¢nding also urges caution before interpolating paleo-pCO2 values between chronologically widely spaced data.

Acknowledgements This research was supported by INSUE. We thank Ocean Drilling Program for permission of sampling cores (hole 534A), and Jean Espitalie¤ (Institut Franc[ais du Pe¤trole) for providing laboratory space to G.D. to carry out carbon analyses. We gratefully acknowledge provision of vertebrate samples by D. Goujet (MNHN, Paris), A. Le¤na and G. Gallio (MHN, Besanc[on), J.H. Delance (University of Burgundy, Dijon), and Museum of Lausanne via M. Weidmann. We are grateful to Andrew Knoll for his critical and constructive remarks on an earlier version of this paper, and Hugh Jenkyns, Robert Gregory for their reviews.[BOYLE]

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