PALEOCEANOGRAPHY, VOL. 18, NO. 2, 1031, doi:10,1029/2002PA000848, 2003 Extreme polar warmth during the Cretaceous greenhouse? Paradox of the late Turonian 8^^0 record at Deep Sea Drilling Project Site 511 Karen L. Bice Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA Brian T. Huber Department of Paleobiology, Smithsonian National Museum of Natural History, Washington, D. C, USA Richard D. Norris Scripps Institution of Oceanography, University of California, San Diego, La Jolla, California, USA Received 13 September 2002; revised 6 January 2003; accepted 15 January 2003; published 7 May 2003. [i] Oxygen isotope data for upper Turonian planktonic foraminifera at Deep Sea Drilling Project Site 511 (Falkland Plateau, 60?S paleolatitude) exhibit an ~2%o excursion to values as low as ?4.66%o (Vienna Peedee belemnite standard; PDB) coincident with the warmest tropical temperature estimates yet obtained for the open ocean. The lowest planktonic foraminifer 6 O values suggest that the upper ocean was as warm as 30-32?C. This is an extraordinary temperature for 60?S latitude but is consistent with temperatures estimated from apparently coeval mollusc ?'^0 from nearby James Ross Island (65?S paleolatitude). Glassy textural preservation, a well-defined depth distribution in Site 511 planktonics, low sediment burial temperature (~32?C), and lack of evidence of highly depleted pore waters argue against diagenesis (even solid state diffusion) as the cause of the very depleted planktonic values. The lack of change in benthic foraminifer 6' O suggests brackish water capping as the mechanism for the low planktonic 6 O values. However, mixing ratio calculations show that the amount of freshwater required to produce a 2%o shift in ambient water would drive a 7 psu decrease in salinity. The abundance and diversity of planktonic foraminifera and nannofossils, high planktonic ibenthic ratios, and the appearance of keeled foraminifera argue against lower-than-normal marine salinities. Isotope calculations and climate models indicate that we cannot call upon more depleted freshwater 6' O to explain this record. Without more late Turonian data, especially from outside the South Atlantic basin, we can currently only speculate on possible causes of this paradoxical record from the core of the Cretaceous greenhouse. INDEX TERMS: 3344 Meteorology and Atmospheric Dynamics: Paleoclimatology; 3337 Meteorology and Atmospheric Dynamics: Numerical modeling and data assimilation; 4267 Oceanography: General: Paleoceanography; KEYWORDS: carbon dioxide, Cretaceous, Turonian, oxygen isotopes, modeling, sahnity Citation: Bice, K. L., B. T. Huber, and R. D. Norris, Extreme polar warmth during the Cretaceous greenhouse? Paradox of the late Turonian 6'**0 record at Deep Sea Drilling Project Site 511, Paleoceanography, 18(2), 1031, doi:10.1029/2002PA000848, 2003. 1. Introduction latitude late Turonian at Deep Sea Drilling Program (DSDP) r T ivjj ^ . j. ^ .? ^ Site 511 (Falkland Plateau, Figure 1) stand out as problem- 121 Mid-Cretaceous upper ocean temperature estimates ,? ,r , , , nnncn _* j ii&r\ i i r; ^ <, ? s /ciSr^N 1 iT 1 1.. ? atic. Huber et al. [19951 reported o O values as low as from oxygen stable isotope (6 0) analyses of planktonic _^_^^,^^^ ^^^^^^^ ^^ ^.^^^^ ^^^^^^ ^^j^^.^^ ^^^ toramimlera are typically 8-20 C warmer than modem sea . .^ , j i w .r ? vr rr,, , ^ c ^ J /cc.T-\ rn- j ?r ? ".i\t\'^-i T- exquisitcly preserved planktonic toramimlera. The data surlace temperatures (SST) \Bice and Norris, 20021. Ter- ^ . i TT i. ^ i ? j- ^ ^. , , *^ ,/ '^ ^, , ? , , ^Z j r. M reported by Huber et al. indicate upper ocean waters as restrial plant assemblages from the high latitudes ot both ^ i,? ^^or^ , ^r.oo i i ..-.. j ? .^i i .. , . , 1 ? T- ^ 2 j warm as 28-30 C at ~60 S paleolatitude m the late hemispheres also indicate warmer-than-modern mean ?._, ,, \ ^ ? i r ^i , , ? ^ ? ? , ^ iT .1 ? ? . 1 Turonian. Today, such temperatures are typical ot the annual and winter minimum temperatures tor this interval ,., ^ n -s- ij ^ j- rrr jo ? i nn^n T- -i ^j. ^ 11 tropical wcstem Paciiic Warm pool and SO sccm cxtraordi- \Herman and Spicer, 19961. Fossil vertebrate assemblages r i ? i i ^-.^ j rj,, ^- ^ ? i ^ ^i r i. i. -j^oxT 1 1 I-.^ J ? i .. Ml 1 nary tor high latitudes. These estimates might therefore be trom ~72 N paleolatitude are consistent with mean annual .r^ ??? ^T ^ r- ? j ?? ^- i^ Z^^or^? ^, r^ ? ^ ? ? rT J . Written otf as an artitact of unrccogmzed diagcnetic altcr- temperatures >14 C m the Turoman-Comacian \Tarduno et ^- i ? j ^ ? ^ ? j i ? ii- i i ; innon ^r ^ J .. ^- ? A r^ ^ ation Or uncxplamcd extreme isotopic depletion ot local ai, 19981. Yet, among data supporting a warm mid-Creta- ^ ,, '^ .?., ^ , r^rvA-n ^i .. J _,, , ^ ? ^ ^- ? c .11 1 ? 1 water. However, Wilson et al. [2002] recently presented ceous earth, paleotemperature estimates trom the high- , ? i , i .^ ? ciSr-. i ? ^\. -^u / *^ ^ ? tropical planktonic o O values consistent with temperatures 6-8?C warmer than modem tropical SSTs from DSDP Site Copyright 2003 by the American Geophysical Union. 144 on Demerara Rise (~2-5?N paleolatitude. Figure 1) in 0883-8305/03/2002PA000848$l2.00 a late Turonian sequence that is broadly coeval with the 9- 1 9-2 BICE ET AL,: PARADOX OF LATE TURONIAN AT SITE 511 Figure 1. South Atlantic Late Cretaceous paleogeographic reconstruction showing the sites discussed in the text. The 92 Ma reconstruction is from the PLATES Project of the University of Texas Institute for Gephysics [Lawver et al, 2001]. KP, Kerguelen Plateau; EANT, East Antarctica; DSDP, Deep Sea Drilling Project. anomalous Site 511 record. Recognition of a late Turonian hyperthermal period in the tropics provides motivation to re-examine the late Turonian sequence on Falkland Plateau. Here we present new Site 511 isotope data from planktonic and benthic foraminifera and the possible implications of the record. 2. Material and Methods [3] DSDP Site 511 (51?00.28'S, 46?58.30'W) is located on continental crust on Falkland Plateau at 2600 m water depth. To obtain a detailed record across the negative 6 O shift reported for Turonian sediments by Huber et al. [1995], we sampled at 5-75 cm intervals in core 47 and the upper part of core 48, which were recovered from between 404.31 and 413.83 m below seafloor (mbsf). The sediments are unconsolidated, primarily grey to dark grey clay stone with a calcium carbonate content ranging from 10 to 23% [Ludwig et al, 1983; Huber et al, 1995]. Sediments in core 47-1 and upper 47-2 are carbonaceous and tended to yield no foraminifera or too few individuals for stable isotope measurements. Below 47-2, 33 cm, the planktonic foraminiferal and calcareous nannofossil assemblages include cosmopolitan species that are associated with nor- mal, open marine depositional settings, an environmental interpretation supported by the sedimentary trace element geochemistry [Varentsov, 1983]. [4] A late Turonian age is assigned to these samples on the basis of co-occurrence of the calcareous nannofossil Kamptnerius magnificus and the planktonic foraminifera Marginotruncana marginata (= Globotruncana bulloides of Huber et al. [1995]) and Praeglobotruncana stephani in the upper 40 cm of section 511-48-1 [Wise, 1983; this study]. In Tethyan sections the first occurrences of K. magnificus and M. marginata are recorded in the late Turonian [Roth, 1978; Premoli Silva and Sliter, 1994] and the last occurrence of P. stephani in the late middle Turonian [Premoli Silva and Sliter, 1994]. Strontium isotope ratios are also consistent with a late Turonian age (below, section 4). [5] The location of the Cenomanian/Turonian boundary below late Turonian sample 48-1, 76 cm (414.26 mbsf) is uncertain because of an 11 -m sequence of red clay stone that contains non- diagnostic agglutinated foraminifera and mollusk and echinoderm fragments and is barren of calcareous foraminifera and nannofossils. The absence of calcareous microfossils may be due to a shallow calcite compensation depth [Wise, 1983]. Planktonic foraminifera reappear farther downcore, in sample 49-2, 25 cm, where they are rare relative to benthic foraminifera, but their abundance increases to 95% of the foraminiferal assem- blage by sample 49-5, 25 cm. Planktonic foraminifera diagnostic of the late Albian, including Rotalipora ticinensis and Rotalipora cf. appenninica, occur 40 cm below this sample. No biomarker species have been recor- ded between this level and the upper Turonian identified in core 48-1. [6] Samples were dried, soaked in distilled water with Calgon, and washed through a 38 p,m sieve. Disaggrega- tion of clay required drying/washing the samples twice. The samples below 405 mbsf contain abundant, small planktonic foraminifera and rare benthic foraminifera. Carbon and oxygen stable isotope measurements were made on monospecific samples of the planktonic foramin- ifera Globigerinelloides volutus, Heterohelix moremani, and Hedbergella sp. from the 106-180 |im fraction, and Whiteinella baltica, Archaeoglobigerina bosquensis, Mar- ginotruncana marginata, and Praeglobotruncana stephani Irom the 212-300 |im fraction. Each sample contained 30- 100 individual planktonic tests in order to assure adequate sample mass, with the exception of M. marginata and P. stephani samples, which contained 3-8 of these larger individuals. Measurements were also made of the benthic foraminifera Gyroidinoides globosus, Berthelina spp. and Dentalina sp., as well as groups of mixed benthics and shallow infaunal benthic specimens. Stable isotope analyses were performed at Woods Hole Oceanographic Institution on a Micromass PRISM-II spectrometer with common acid bath. Precision of the measurements is ?0.03%o for carbon and ?0.05%o for oxygen. All carbonate isotope data here are reported relative to the Vienna Peedee belemnite (PDB) isotope standard. 3. Results 3.1. Foraminiferal Preservation and Calcareous Plankton Assemblages [7] Calcareous benthic foraminifera typically have glassy, translucent tests throughout core 47. Glassy specimens of planktic foraminifera can also be found throughout the core when viewed as dry specimens, but constitute the majority BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 9-3 Figure 2. Scanning electron micrographs of foraminifera from Deep Sea Drilling Project Site 511. (a)- (c) Wall microstructure, note primary microlayering and absence of secondary calcite on inner and outer shell wall. Scale bars = 5 (xm. (a) Planktonic Whiteinella baltica, 511-47-3, lli-lS cm. (b) Benthic Gyroidinoides globosus, 511-47-5-13.5-15.5 cm (c) Benthic Gyroidinoides globosus, 511-48-1-24-27 cm. (d)-(g) Whole specimens from 511-47-5, 23-25 cm. (d) Schackoina cenomana, scale = 100 |xm. (e) Whiteinella baltica, scale =10 \an. (f) Laeviheterohelix, scale = 10 \an. (g) Globigerinelloides sp., scale = 100 ixm. (>85%) of specimens only between 47-6, 72 cm and 47-3, 23 cm. Above and below this interval in core 47, specimens display excellent preservation, but have mostly opaque skeletons. In section 48-1, foraminifera are texturally well preserved but are typically opaque and in many cases fragmentary. Preservation of glassy tests, which is very rare for foraminifera of this age, can be attributed to the moderate burial depth and low porosity and permeability of the enclosing clay-rich sediments. Primary foraminifer wall microstructure preservation is excellent throughout (Figure 2), with distinct shell layering and wall pores clearly visible even in the material that could not be described as "glassy" in dry specimens. There is no visible evidence of recrystallization or addition of secondary calcite inside or outside the shell walls in any foraminifera samples. Calca- reous narmofossils from all samples in the late Turonian interval were rated as moderately preserved by Wise [1983] because of etching as a result of minor dissolution. [8] The foraminiferal abundance and species diversity follows similar trends to skeletal transparency, with the greatest species richness (up to 16 species) of the entire Cretaceous sequence, the greatest individual abundance, and the highest percentage of keeled planktonic species in sections 47-4 to the top of 47-6. Such dominance of the foraminifer assemblages by planktonic specimens suggests the sediments were deposited below 1000 m water depth on the basis of analogy with modem planktonicibenthic fora- miniferal ratios [Gibson, 1989]. The increased species diversity and presence of the keeled species in the late Turonian are consistent with warming and/or increased vertical stratification of a normal marine, oligotrophic surface mixed layer. Calcareous nannofossils from this interval similarly indicate their growth in relatively warm, normal marine and oligotrophic surface waters during the late Turonian (S. Wise, personal communication, 2002). [9] Planktic foraminifera are almost entirely absent above sample 47-1, 106 cm and are present almost exclusively as atypically small specimens in the lower part of section 47-1 and throughout 47-2. In section 48-1, planktic foraminifera vary between 50 and 95% of the total assemblage and are lower in diversity compared with those in much of core 47. 3.2. Stable Isotopes [10] The new isotope data are consistent with values reported by Huber et al. [1995]. Both sets of data are shown in Figure 3 and our new data are listed in Table 1. There is an abrupt decrease upsection in planktonic 6^^0 between 412.22 and 412.14 mbsf. Below this level, 9-4 BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 III LU Ll. LU 7" OH I w CO n o 1- CD < N o _j ^ TEMPERATURE {?C) 12 16 20 24 28 32 ~ 405 - 406 - CD I"* I CD CD <-> o ?> 001 407 - - 408 - -_- 409 - I-:^ 410 - ^^ 411 - 412 - 413 - I I I I I I I I I I I I Benthic T(?C) 16 20 r~rn I to ? ? ?? -> ft ?^^?^-?? ???(P 7 c m _i_ < m g SI j? 1 9- N"! 1 < CD ?? _i DC I 1 414 429 4301- J_L ll^L I I I I I I I I I I T ? Vi V ?^ ? 940 ? ? r ^ 111 ^^i*N^i _ 1 1 1 1 1 1 1 1 1 l_ .?^ ^^ - - ? T 1 1 1 I I I I I I I I I I ? V -1 -2 -3 -4 5^^0 (PDB) -1 0 1 2 8^^C (PDB) PLANKTONIC A M. marginata ^ w. baltica BENTHIC ? G. volutus n H. moremani n Praeglob. spp. ? Shallow infaunal ? Mixed T Hedbergella sp. O 'A- bosquensis H. planispira A Deep infaunal Figure 3. Oxygen and carbon isotope data for DSDP Site 511, cores 47-1 through 49-5. Planktonic foraminifer temperature calculations assume 6^ = ? 1.5%o SMOW. Benthic foraminifer temperature calculations assume 6^ at 1000 m water depth = ? 1.0%o. The horizontal line indicates the position of a subtle lithologic change observed in the archive half of core 511-47-6 at 412.18 mbsf (see text). Other such lithologic changes occur in core 47 but their positions are not noted here. The lithologies are clay stone with rare zeolites (Turonian) and calcareous ooze (Albian). Black fdl represents core recovered. The core number is shown in the core column. See color version of this figure at back of this issue. Turonian 6'**0 averages 2.2%o (PDB) with low (<0.4%o) variance and late Albian 6'^0 values average ? 1.4%o (PDB). Above this level, planktonic oxygen values decrease abruptly by ~2%o and show good separation. Margin- otruncana marginata, H. moremani and A. bosquensis generally exhibit the lowest values and Hedbergella and G. volutus have the highest values. Two samples of W. baltica exhibit values below ?4.00%o, including the second lowest 6'^0 value (?4.47%o) we measured, but otherwise W. baltica varies within the spread of planktonic 6 O values. M. marginata yielded the lowest 8'^0 measurement (?4.66%o). With the exception of an anomalous (in 6^^0 BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 9-5 Table 1. Oxygen and Carbon Isotopic Values for Planktonic and Benthic Foraminifera From DSDP Site 511 Core Interval, cm Depth, mbsf Taxon 6"C, %o 6"*0, %o 47-1 106-108 405.06 Hedhergella sp. 0.746 -1.721 47-2 106-108 406.56 Hedbergella sp. -0.048 -1.768 47-2 106-108 406.56 Heterohelix moremani -0.483 -2.467 47-2 120-122 406.70 Hedhergella sp. 1.826 -2.169 47-2 120-122 406.70 mixed benthics 0.176 -1.222 47-3 34-36 407.34 Heterohelix moremani 0.758 -3.106 47-3 34-36 407.34 Whiteinella baltica 1.453 -2.485 47-3 34-36 407.34 mixed benthics 0.457 -1.459 47-3 105-107 408.05 Globigerinelloides volutus 1.592 -2.224 47-3 105-107 408.05 Hedbergella sp. 1.294 -2.205 47-3 105-107 408.05 Whiteinella baltica 1.586 -2.763 47-3 105-107 408.05 Heterohelix moremani 1.312 -2.945 47-4 30-32 408.80 Whiteinella baltica 2.104 -3.578 47-4 106-108 409.56 Globigerinelloides volutus 1.276 -2.816 47-4 106-108 409.56 Hedbergella sp. 1.296 -2.906 47-4 106-108 409.56 Whiteinella baltica 1.528 -3.073 47-4 106-108 409.56 Heterohelix moremani 1.161 -3.862 47-4 106-108 409.56 Archaeoglobigerina bosquensis 1.757 -3.374 47-4 106-108 409.56 mixed benthics -0.904 -1.682 47-5 33-36 410.33 Globigerinelloides volutus 1.499 -3.031 47-5 33-36 410.33 Hedbergella sp. 1.632 -3.120 47-5 33-36 410.33 Whiteinella baltica 1.762 -3.466 47-5 33-36 410.33 Heterohelix moremani 1.062 -3.802 47-5 33-36 410.33 Archaeoglobigerina bosquensis 1.737 -3.921 47-5 33-36 410.33 Marginotruncana marginata 1.219 -4.115 47-5 33-36 410.33 Gyroidinoides globosus -0.655 -1.932 47-5 33-36 410.33 mixed benthics -0.533 -1.675 47-5 13.5-15.5 410.14 Whiteinella baltica 1.706 -4.078 47-5 13.5-15.5 410.14 Marginotruncana marginata 0.849 -3.980 47-5 13.5-15.5 410.14 Gyroidinoides globosus -0.256 -1.477 47-5 72-74 410.72 Whiteinella baltica 1.779 -4.014 47-5 72-74 410.72 Archaeoglobigerina bosquensis 1.627 -3.876 47-5 72-74 410.72 Marginotruncana marginata 1.221 -4.154 47-5 72-74 410.72 Gyroidinoides globosus -0.338 -1.653 47-5 109-111 411.09 Globigerinelloides volutus 1.218 -2.938 47-5 109-111 411.09 Hedbergella sp. 1.369 -3.164 47-5 109-111 411.09 Whiteinella baltica 1.930 -4.085 47-5 109-111 411.09 Heterohelix moremani 0.995 -4.164 47-5 109-111 411.09 Archaeoglobigerina bosquensis 1.707 -4.334 47-5 109-111 411.09 Marginotruncana marginata 0.941 -3.610 47-5 124-126 411.24 Whiteinella baltica 1.635 -4.473 47-5 124-126 411.24 Archaeoglobigerina bosquensis 1.429 -4.089 47-5 124-126 411.24 Berthelina sp. 0.598 -1.585 47-5 124-126 411.24 mixed benthics -0.920 -1.476 47-6 18-20 411.68 Marginotruncana marginata 1.673 -4.664 47-6 18-20 411.68 Dentalina sp. -0.789 -1.753 47-6 18-20 411.68 shallow infaunal benthics -0.318 -1.612 47-6 31-33 411.81 Globigerinelloides volutus 1.522 -2.369 47-6 31-33 411.81 Hedbergella sp. 1.764 -2.908 47-6 31-33 411.81 Whiteinella baltica 2.061 -3.677 47-6 31-33 411.81 Heterohelix moremani 1.107 -3.277 47-6 31-33 411.81 Archaeoglobigerina bosquensis 2.081 -3.589 47-6 31-33 411.81 Marginotruncana marginata 1.690 -4.240 47-6 31-33 411.81 Marginotruncana marginata 1.849 -4.191 47-6 38-40 411.88 mixed benthics -0.111 -1.542 47-6 64.5-66.5 412.14 Whiteinella baltica 1.647 -4.282 47-6 64.5-66.5 412.14 Marginotruncana marginata 1.172 -3.584 47-6 64.5-66.5 412.14 Dentalina sp. -1.807 -2.246 47-6 72-74 412.22 Hedbergella sp. 0.835 -2.348 47-6 72-74 412.22 Praeglobotruncana delrioensis -0.006 -2.236 47-6 72-74 412.22 trocospiral benthics -0.206 -1.664 47-6 72-74 412.22 deeper infaunal benthics -1.129 -1.830 47-6 98-100 412.48 Praeglobotruncana stephani 0.275 -2.390 47-6 98-100 412.48 Berthelina sp. 0.613 -1.334 47-6 107.5-109.5 412.58 shallow infaunal benthics 0.400 -1.742 47-6 105-107 412.55 Globigerinelloides volutus 1.501 -2.146 47-6 105-107 412.55 Hedbergella sp. 1.358 -2.244 47-6 126-128 412.76 Berthelina sp. 0.333 -1.381 48-1 24-27 413.74 Berthelina sp. 0.631 -1.475 48-1 33-35 413.83 Whiteinella baltica 1.154 -2.327 48-1 33-35 413.83 Whiteinella baltica 1.238 -2.371 9-6 BICE ET AL,: PARADOX OF LATE TURONIAN AT SITE 511 Table 1. (continued) Core Interval, cm Depth, mbsf Taxon 6"C, %o 6'**0, %o 48-1 33-35 413.83 Hedhergella sp. planispira 1.324 -2.194 48-1 33-35 413.83 Praeglobotruncana stephani 0.339 -2.075 48-1 33-35 413.83 Praeglobotruncana stephani 0.295 -2.028 48-1 33-35 413.83 mixed benthics -0.087 -1.582 48-1 33-35 413.83 deeper infaunal benthics 0.013 -1.542 49-5 65.5-67.5 429.36 Hedhergella sp. planispira 1.511 -0.811 49-5 65.5-67.5 429.36 Hedhergella sp. portsdownensis 1.370 -1.476 49-5 65.5-67.5 429.36 Berthelina sp. 1.086 -0.558 49-5 65.5-67.5 429.36 Berthelina sp. 1.376 -0.448 49-5 103-105 430.04 Hedhergella yezoana 2.784 -1.457 49-5 103-105 430.04 Berthelina sp. 1.704 -0.555 and 8 C) Dentalina measurement at 412.14 mbsf, benthic values show little change through the study interval. [ii] Cross plots of carbon and oxygen isotope data provide an indication of planktic foramiferal depth ecology. Within core 47, G. volutus, H. moremani, and Hedhergella spp. consistently plot as lower mixed layer species. A. bosquensis and W. baltica appear to be upper mixed layer species, and M. marginata occupies a mid to upper mixed layer position. The separation of these species from co- occurring benthics is consistenly large through the interval. [12] The abrupt change in 6 O that occurs between 412.22-412.14 mbsf suggests the possibility of an uncon- formity. There is a subtle lithologic change at 412.18 mbsf where sediment color changes from a medium dark gray below to lighter gray above. Carbonate content, foramini- fer abundance and planktonic species diversity increase above this point. XRD analysis of six samples taken within 2 m above and below the 6 O shift show no significant change in mineralogy of the bulk sediments. Sediment color alternates between darker and lighter gray throughout most of core 47-6, but the change is often more gradual than that at 412.18 mbsf. If the subtle lithologic change marks a hiatus, the amount of missing time is entirely unconstrained. 3.3. Temperature Estimates [13] To calculate upper ocean temperatures from plank- tonic foraminiferal 6 O, we use the equation of Erez and Luz [1983]. Ambient water 8'^0 value (6w) is estimated (relative to Standard Mean Ocean Water, SMOW) and is converted to the PDB scale by subtracting 0.27%o [Hut, 1987]. We first use the standard approach to estimating 6w, which assumes an "ice-free" mean ocean 6^ of ? 1.0%o SMOW [Shackleton and Kennett, 1975] and includes an adjustment that takes into account average latitudinal variations in evaporation and precipitation controls on 6'^0 by analogy with the modem ocean [e.g., Zachos et al, 1994]. At a latitude of 60?, this approach yields 6^ = -1.2%o (SMOW) and paleotemperatures of 19-2rC below 412.18 mbsf Above this depth, calculated temperatures are as warm as 30-32?C for shallower dwelling M. marginata, W. baltica, and A. bosquensis, and 23-25?C for deeper dwelling planktonic foraminifer species. More conserva- tively, we can assume that surface 8w at 511 was as depleted relative to the global mean as the lowest water 8'^0 in the Geochemical Ocean Sections Study (GEOSECS) [1987] data set poleward of 50?S latitude. This value (?0.48%o SMOW) occurs at 1 m water depth at station 82 in the austral summer, near the South Sandwich Islands, at 56.3?S latitude. Assuming a local 6^ of ? 1.5%o (~?0.48 + ? 1.0%o) decreases the temperature estimates given above by 1.4?C. The more conservative 6w estimate results in maximum upper ocean temperatures of 28-30?C for shallower dwelling species and 22-24?C for deeper dwell- ing planktonic foraminifera (Figure 3). [14] Assuming a range of possible bottom 6w values between average "ice-free" seawater (?1.0%o SMOW) and the conservative upper ocean water estimate (?1.5%o SMOW), we estimate local bottom water temperatures (>1000 m) to have been 16-20?C throughout the study interval. The data indicate a local vertical temperature gradient of only ~4?C below 412.18 mbsf and as great as 12?C above this point. 4. Comparison With Other Late Turonian Data [15] We compared our data from DSDP Site 511 with age equivalent records from other sites where synchroneity could be determined through a combination of biostratig- raphy and strontium (Sr) isotope ratios. Unfortunately, few age equivalent records exist. Sr isotope data are available for foraminifera reported by//M^er e^ a/. [1995] forSite511 and equatorial Atlantic Site 144 [Wilson et al., 2002], as well as new data for Site 511 (Table 2). After renormaliza- tion to NIST 987 = 0.710248 (as required to compare the data to the Sr reference curve oi McArthur et al. [2001]), these records are age equivalent within an interval of <600 kyr in the late Turonian. Samples for DSDP Site 511" Depth, Standard Sample mbsf 87Sr/86Sr'' Error, % Type 47-5, 13.5-15.5 cm 410.14 0.707314 0.0006 Foraminifera 47-6, 18-20 cm 411.68 0.707323 0.0007 Foraminifera 48-1, 14-18 cm 413.21 0.707408 0.0006 Foraminifera 48-1, 33-35 cm 413.83 0.707377 0.0007 Foraminifera 48-2, 84-86 cm 415.84 0.707379 0.0006 Inoceramid prisms "See Huher et al. [1995] and Bralower et al. [1997] for additional Cretaceous Sr values for this site. All samples analyzed by Paul D. Fullagar at the University of North Carolina. 87Sr/86Sr values are corrected using the difference between the laboratory average of 0.710270 for the NIST 987 carbonate standard and the accepted value of 0.710250 for NIST 987. BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 9-7 [i6] Ditchfield et al. [1994] report oyster 6'^0 values of -1.75, -4.64 and -6.09%o PDB from the mid Cretaceous on James Ross Island (JRI, paleolatitude ~65?S, Figure 1). In paleotemperature calculations, Ditchfield et al. exclude the two lower values, apparently on the basis of textural preservation or the isotopic values themselves [Ditchfield et al, 1994, p. 86]. However, based on iron and manganese concentrations reported by McArthur et al. [2000] for these samples and the criteria for unaltered material defined by Ditchfield et al. (<500 ppm Fe, <2G0 ppm Mn), only one of the oyster 8 O values reported from JRI should have been used in paleotemperature calculations: that with an oxygen isotope value of -6.09%o PDB. On the basis of SEM examination, McArthur et al. [2000] describe this sample (D.8228.113) as flaky, translucent calcite characteristic of pristine oyster carbonate. [n] McArthur et al. [2000] showed that, based on ^^Sr/^'^Sr, JRI sample D.8228.113 dates to the late Turonian strontium minimum at ~89.45-89.74 Ma. Normalizing all the Sr data from Sites 144, Site 511, and James Ross Island to NIST 987 = 0.710248 [McArthur et al, 2001] shows that the three data sets date to within a 1.5 myr interval of the late Turonian. We note, however, that high-resolution mid- Cretaceous records [e.g., Wilson and Norris, 2001] show substantial climate variability within such a time window. We are therefore severely limited it what we can say about changes in late Turonian global climate given the data available for comparison. [18] Wilson et al. [2002] have inferred late Turonian sea surface temperatures of 30-33?C from their data suggesting that the tropical Atlantic had temperatures similar to or slightly warmer than maximum temperature estimates from Site 511 (Figure 4). Paleotemperatures implied by the oyster data on James Ross Island are hard to determine precisely since oysters can live in much lower-salinity waters (influenced by local runoff) than is generally believed to be true for planktonic foraminifera. Here we make the liberal ad hoc assumption that the oysters grew in waters 2.0%o more depleted than open ocean waters at equivalent latitude. This is a generous allowance for depleted water given that the JRI faunal assemblage is interpreted as indicative of normal marine salinities [Ditchfield et al, 1994]. If the mean ocean 6^ was -1.0%o SMOW, our assumptions yield a 8w of ?3.5%o for the oyster sample. The paleotemperature equation for mollusc carbonate [Anderson and Arthur, 1983] then indicates that shelf waters around James Ross Island were about 27?C in the late Turonian. (A smaller 6w adjustment for runoff would result in a higher temperature estimate.) Hence the oyster data are in broad agreement with the temperature estimates derived above for DSDP Site 511 and reinforce an interpretation of a small latitudinal temperature gradient during the core of the Cretaceous greenhouse climate (Figure 4). [19] Somewhat lower-temperature estimates result for all these sites if we take into account the long term trend in mean 8w due to the imbalance between '^O sources, primarily high-temperature seafloor alteration, and sinks such as continental weathering, low-temperature seafloor alteration, and water fixation in the upper oceanic crust and in clay minerals. Wallmann [2001] estimated that the mean 36 - 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 32 - 'h \ b?. - 24 - v/ '?^?^ ~ 20 - J V \ ~ ^?? \ \- 16 ? \ \ \ ? ^ "r 12 - ^ ? ^ 7500 7500 ppm ppm QFACT=25 QFACT=20 \S ppm QFACT=15 \\ 8 - " " ? 4500 ppm QFACT=12 1 1 1 I. 11 A 1 1 1 1 1 1 803 60S 40S 20s 20N 40N 60N SON Latitude ? G. valutas ? W. baltica V Hedbergella sp. D H. moremani ^ oyster, James Ross I. ? M. marginata O A. bosquensis Figure 4. Curves show mean January (austral summer) sea surface temperature profiles from four GENESIS version 2.0 AGCM experiments. Symbols show estimated upper ocean temperatures from 6 O of approximately coeval late Turonian samples. Isotopic temperature calculations assume the following 6^ values: Site 144 foraminifera (5?N lat.) 6w = -0.25%o SMOW; Site 511 foraminifera (60?S lat.) 6w = ? 1.5%o; James Ross Island mollusc (65?S lat.) 6w = ?3.5%o (see text). Symbol shapes for foraminiferal species are the same as those used in Figure 3. All Site 144 estimates are fi-om H. moremani (squares) and Hedbergella sp. (inverted triangles). Site 144 data are from Wilson et al. [2002]. The James Ross Island datum is from Ditchfield et al. [1994]. See color version of this figure at back of this issue. ocean 6^ at 100 Ma would have been depleted ~0.25%o beyond the ice-free value, yielding an ocean mean of -1.25%o SMOW. This -0.25%o adjustment would decrease all isotopic paleotemperature estimates shown in Figure 4 by ~1.2?C. [20] If the Site 511 late Turonian data record an interval of extreme polar warmth, the most likely mechanism is a substantial increase in atmospheric greenhouse gas concen- trations and increased atmospheric poleward heat transport. Using annual mean conditions predicted from the GENESIS version 2.0 atmospheric general circulation model (AGCM), Bice and Norris [2002] determined that, if the anomolously high temperatures implied by the Site 511 late Turonian data are ignored, maximum late Albian-Turonian temperatures can be reproduced with 4500 ppm CO2 and the consequent water vapor feedbacks. When the maximum temperatures estimated from the Site 511 data need to be matched, 6500-7500 ppm CO2 is required in the AGCM. Poleward heat transport must be increased by at least 50% Irom the "background" warm mid-Cretaceous conditions. Mean zonal January (austral summer) temperature profiles Irom these models are shown in Figure 4. 9-8 BICE ET AL,: PARADOX OF LATE TURONIAN AT SITE 511 [21] Any rise in high-latitude surface temperature occur- ring during a Turonian greenhouse climate should also affect temperatures at the sites of bottom water formation and increase temperatures estimated from benthic foramin- ifera. However, there is no significant change in bottom water temperature estimates at Site 511 associated with the interval of low planktonic 6 O values (Figure 3). Indeed, when the GFDL Modular Ocean Model version 2 is forced with the mean annual atmospheric output from AGCM runs with 6500-7500 ppm CO2, the bottom water temperature predicted is ~28?C, which is ~10?C warmer than bottom temperatures estimated from the benthic data. High-latitude winter SSTs predicted by the AGCM with 6500-7500 ppm are 23-25?C. Therefore, if the ocean model is forced by seasonally varying surface temperatures and bottom water formed only during winter months, bottom waters would still overpredict the isotopic paleotemperatures by 6-8?C. [22] We can produce a reasonable match to Site 511 benthic paleotemperatures in model runs using 4500 ppm CO2. Although experiments with 4500 ppm CO2 under- predict the warmest estimated surface ocean temperatures (Figure 4), the mismatch is not much greater than the error in the isotopic paleotemperature calculation. Assuming that the general circulation models are reliable tools for studying greenhouse climates, the models suggest that, if the Site 511 late Turonian data are a record of surface warming alone, then carbon dioxide concentrations would have likely been 4500 ppm or greater. Carbon dioxide concentrations of this magnitude are inferred from leaf stomatal index measure- ments of Cenomanian and Turonian Ginkgo fossils [Retallack, 2001] but are higher than values derived Irom other mid-Cretaceous proxies [Bice and Norris, 2002]. 5. Alternative Explanations for the Site 511 8^*0 Data 5.1. Possible Diagenesis [23] It appears highly unlikely that the low 6 O plank- tonic foraminifer values result from diagenetic alteration. The significant separation among the planktonics and between planktonics and benthics argues against the low 6'^0 being caused by postdepositional calcification or alteration. There is no textural evidence for infilling, encrustation or recrystallization. The high quality of plank- tonic foraminiferal preservation is also refiected by the abundant occurrence above 412.18 mbsf of delicate biserial foraminifera, which are rare lower in the section. [24] Authigenic clay formation, which can promote low pore water 6'^0 [Lawrence et al, 1975] appears to be a unlikely mechanism for alteration here. Robert and Maillot [1983] report no evidence for authigenic clays and the absence of evidence of burial, volcanic or organic diagenesis. The predominant clay mineral in the Turonian sequence is pedogenic smectite, interpreted as having been derived from deep soils in downstream, low-relief areas of a continental drainage basin in an area of constantly warm, wet climate [Robert and Maillot, 1983; Timofeev et al, 1983]. [25] Huelandite group zeolites are common in the Santo- nian through Maastrichtian section at Site 511, but are rare below this level, including the interval we studied [Varent- sov et al, 19S3; Huber et al., 1995]. Varentsov et al. [1983] report that the highest content of fine basaltic volcaniclastic sediments that have been altered to huelandite occurs in the Campanian-Maastrichtian interval, yet foraminiferal 6 O values for this interval are quite enriched [Huber et al, 1995]. This argues against diagenesis of volcaniclastic material as a source of light oxygen in the late Turonian interval. [26] DSDP Site 511 sits on continental crust, so pore water interaction with oceanic crust basalts is an unlikely mechanism for the introduction of depleted oxygen. Site 511 sediment temperature measurements made at 0, 52 and 113 mbsf indicate a local temperature gradient of 74?C km~^ [Langseth and Lugwig, 1983] and indicate a sediment temperature at 412 mbsf of only ~32?C. Cretaceous sediments crop out in various places on Falkland Plateau as a result of Maastrichtian and Eocene stripping [Ludwig, 1983]. It is therefore possible that the Turonian sequence was more deeply buried in the past. However, the relatively good preservation of foraminifera in the chalkier parts of the Cretaceous sequence suggests that they were never very deeply buried, and Robert and Maillot [1983] note the absence of any textural evidence for burial diagenesis in these sediments. Significant alteration of the primary 6'^0 signal through solid state diffusion of oxygen isotopes at temperatures below 50?C is unlikely, even if we assume that pore waters had somehow become strongly O-depleted at some point following burial [Anderson, 1969; Kronenberg et al, 1984]. 5.2. Upper Ocean 6? Decrease Through Increased Freshwater Flux [27] The fact that benthic 6^^0 is unchanged through the Site 511 Turonian interval suggests that the planktonic foraminifera record a period during which local water became more O-depleted and that bottom waters were largely unaffected by any coincident climate or oceano- graphic change. Assuming no accompanying temperature change, the data indicate that waters over Falkland Plateau became ~2%o more depleted relative to the global mean. Two broad mechanisms can decrease local water 6'^0: an increased flux of O-depleted freshwater and a decrease in the 6 O of local runoff and precipitation. These two changes need not be mutually exclusive, but they can properly be considered separately. [28] If we view the modem Southern Ocean surface water average 6w (~?0.3%o) as resulting from a simple mixing of average seawater (0%o) and depleted precipitation and run- off (~?10%o) [Rozanski et al, 1993], then the fraction of upper ocean water contributed by precipitation and runoff today is ~3%. In order to decrease local 8^ by 2%o without invoking a change in the oxygen isotopic composition of precipitation/runoff, the freshwater contribution would have to increase to 23% (or to 20-22% if the Cretaceous mean ocean h^ was ?1.0 to 1.25%o). [29] An increase in high-latitude precipitation rate is an expected consequence of tropical warming [Manabe, 1996] and a stronger hydrologic cycle. Hence some increase in high-latitude precipitation rate might be expected in BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 9-9 response to the late Turanian trapical warmth [Wilson et al, 2002]. However, an increase to a 23% freshwater contribu- tion implies a 6-7 psu decrease in local salinity, again, using a simple mixing ratio. The observed maximum in planktonic abundance and diversity in Site 511 in the interval with the lowest 6 O argues against such a salinity decrease. Leckie [1987] showed that keeled foraminifera (e.g., M. marginata) are most abundant in open ocean settings, suggesting a preference for relatively normal salinities. While we cannot rule out some decrease in salinity, the increase in faunal diversity, high planktonic :- benthic ratios and abundance of keeled forms where planktonic foraminiferal 6 O becomes very low are more consistent with a change from abnormal salinity conditions below 412.18 mbsf to more normal marine salinity above. At 412.18 mbsf, we observe a subtle increase in carbonate content and consequent decrease in clay percentage upsection, into the interval of depleted 8 O. A substantial increase in continental runoff to the basin might be reflected in an increase in detrital clay flux, but the opposite change occurs. Given the lack of any change consistent with decreased salinity or increased runoff, we do not believe the Site 511 6 O record can be explained by a substantially increased freshwater flux. 5.3. Upper Ocean 8? Decrease Through Decreased Freshwater 5*^0 [30] Is it possible that surface waters could have become depleted by 2%o relative to the global mean because of a decrease in the 6'^0 of freshwater? Using again the esti- mated modem Southern Ocean freshwater mixing ratio of 3%, in order to decrease local 6^ by 2%o using only the composition of freshwater, local precipitation/runoff 6 O would have had to decrease by 58%o or more. [31] Modem precipitation reaches a minimum 6^^0 of -50%o SMOW over the South Pole [Dansgaard, 1964]. Experiments with atmospheric general circulation models that include water isotopes tracers indicate that high-latitude precipitation would be more enriched in O in a warmer climate. Using modem boundary conditions, Jouzel et al. [2000] show that the climatic changes accompanying an increase in CO2 from 315 to 630 ppm CO2 would increase high-latitude precipitation 8 O by 3-4%o in the north and by 3-7%o in the south. Recent experiments with an isotope-capable version of GENESIS and mid-Cretaceous boundary conditions [Bice et al, 2002] are in agreement with the Jouzel et al. model prediction of isotopically enriched high-latitude precipitation. In the GENESIS model [Mathieu et al, 2002] with 4500 ppm CO2, the predicted 6 O of precipitation over Falkland Plateau varies between ?5 and ?7%o SMOW (summer-winter). If the model predictions that a warmer world has a lower gradient in precipitation 6'^0 are correct, then we cannot invoke more depleted high-latitude freshwater to explain the Site 511 record. [32] Where do waters 2%o lower than the global mean occur today? Ocean water 6'^0 ?2.0%o or lower is recorded in the North Atlantic near 65 ?N, where the East Greenland Current supplies ^^0-depleted Arctic water [Azetsu-Scott and Tan, 1997] to water depths less than 50 m. Water 6 O values lower than ?2%o have also been measured in fiords, within ~100 km of the melting terminus of a tidewater glacier [Azetsu-Scott and Tan, 1997]. In one study near the Antarctic West Ice Shelf at 63-65?S, Frew et al. [1995] found no water 8 O value lower than ?0.5%o, consistent with all GEOSECS [1987] Southern Ocean measurements. It therefore appears that the modem analogs to the 6w required (if the late Turonian data are to be explained by 6w change) are associated with Arctic ocean outflow and tidewater glacier melting. However, the appearance of local glacial ranoff in the Turonian is not supported by any lithologic changes that might be expected in this scenario (e.g., increased clastic flux), and benthic foraminifer 8 O do not record coincident changes in temperature or 8w, although such changes might be expected given a glacial mechanism. 6. Concluding Remarks [33] Our new results, combined with those of Huber et al [1995], clearly show that subpolar planktonic foraminifera over Falkland Plateau became strongly '^O depleted in the late Turonian, about the time of the warmest tropical temperatures thus far estimated for the open ocean [Wilson et al, 2002]. We do not yet have a satisfactory explanation for the extremely depleted planktonic foraminifer 8 O values above 412.18 mbsf at Site 511. Diagenesis appears unlikely as a mechanism given the good separation among planktonic S'^O values, lack of evidence for '^O-depleted pore fluids, low burial temperature, and absence of any textural evidence for alteration or secondary calcification. [34] We do not mle out warming: The narmofossil assem- blage changes at 412.18 mbsf have been interpreted as consistent with some warming (S. Wise, personal commu- nication, 2002), and the upper ocean temperatures derived here are consistent with those from 8 O of well-preserved late Turonian bivalves [Ditchfield et al, 1994], if our assumption of ?3.5%o 6^ at James Ross Island is adequate. However, it is difficult to accept tropic-like temperatures at such high latitudes, especially when available benthic foraminifer data do not indicate coincident bottom water warming. Both the planktonic foraminifera we analyzed at Site 511 and the oysters analyzed by Ditchfield et al. [1994] from James Ross Island could have grown during the austral summer. Climate model experiments require very high atmospheric CO2 concentrations to reproduce the summer temperatures estimated from the two high-latitude data sets (Figure 4). These CO2 amounts are consistent with some estimates of mid-Cretaceous atmospheric CO2 concentra- tions from fossil leaf stomatal index measurements of Ginkgo specimens but are significantly higher than estimates reconstracted from marine organic carbon [Free- man and Hayes, 1992; Retallack, 2001; Bice and Norris, 2002]. However, in order to explain both the planktonic and benthic 6 O records from Site 511 as wholly temperature records we would have to imagine that summer upper ocean temperatures increased locally while temperatures at the sites of deep water formation during winter (the more likely time of deep water formation) did not. No suitable records exist to make this analysis. 9- 10 BICE ET AL,: PARADOX OF LATE TURONIAN AT SITE 511 [35] We also cannot rule out isotopically depleted upper ocean water. However, if the Site 511 record reflects an interval of decreased upper ocean water 6^^0, we need a satisfactory mechanism(s) that would lead to depletion of local waters by ~2%o relative to the global mean. Today, such depletion is associated with polar ocean outflow and tidewater glacier melting, but there is no evidence to support such a mechanism for the late Turonian greenhouse. Iso- tope-capable general circulation models support high-lati- tude precipitation and runoff that is more enriched in ^^O during a warmer climate, not more depleted. Tropical warming can be expected to lead to a stronger hydrologic cycle [Manabe, 1996] with higher precipitation rates at high latitudes and consequent increased runoff. However, whether through direct rainfall or runoff, the increase in freshwater flux to Site 511 required to effect a 2%o decrease in local water would likely have driven a salinity decrease of 6-7 psu. In contrast, the Site 511 record and conventional interpretations of abundance and diversity changes in foraminifera and nannofossils would argue for a change to more normal marine salinities above 412.18 mbsf, not a freshening. [36] The Cretaceous South Atlantic was a smaller, more restricted basin than the modem (Figure 1). We can speculate that the planktonic foraminifer low 6'^0 interval might record isotopic depletion due to increased basin isolation. However, the paleogeography that included a restricted South Atlantic existed long before the late Turonian and yet no record of severe 6w depletion has been identified in the Albian section on Falkland Plateau \Huber et al, 1995; Fassell and Bralower, 1999]. Also, if tectonic change led to severe restriction of the local basin, we might expect a coincident change in bottom water conditions, but none is indicated by the benthic foraminifer isotope records. [37] We can speculate that perhaps Site 511 sat at the juncture of two water masses [e.g., Ciesielski et al, 1977] and that atmosphere and/or ocean circulation changes occurred to bring lower 6^^0, normal salinity waters onto Falkland Plateau. Presumably, such changes might have been related to the tropical warmth noted in the Site 144 record [Wilson et al, 2002]. Two problems exist with a speculative high-latitude water mass mixing source, how- ever. Such water would be expected to be cool and therefore inconsistent with the nannofossil assemblage change (section 3.1) interpreted as consistent with some warming (S. Wise, personal communication, 2002). Also, in order to explain the Site 511 shift, a lower 6 O water mass from high latitudes would have to be at least as depleted as modem Arctic outflow or tidewater glacial drainage. As discussed above (section 5.3), this seems unlikely in the warm mid-Cretaceous. [38] Alternatively, we can speculate that the record is the result of some combination of warming and increased freshwater flux. Any plausible argument must explain upper ocean change without coincident bottom water change at Site 511. We must be careful to not base too much on one data point from James Ross Island, but the most satisfactory explanation for the Site 511 record is one that would also explain extreme '^O depletion (?6.09%o) [Ditchfield et al, 1994] in what has been termed pristine oyster carbonate \McArthur et al, 2000] of approximately equal age. [39] For now, we are left with a paradoxical record in the upper Turonian on Falkland Plateau. Understanding the exact cause of the low 6^^0 interval here will require, as a minimum, coeval upper ocean records from equivalent latitudes in either hemisphere, preferably both within and outside the South Atlantic basin. In addition, more benthic foraminifer records are needed in order to understand what global bottom water and high-latitude changes occurred approximately coeval with the extreme tropical warmth of the late Turonian. [40] Acknowledgments. This material is based upon work supported by the National Science Foundation under grant ATM-0000545, which supports the Partnership in Modeling Earth System History (PSUAVHOI) and grant EAR-9909658. Support was also provided by the Smithsonian Walcott Fund, the Smithsonian Scholarly Studies Fund, the Ralph E. Hall Endowed Fund for Innovative Research, and the Andrew W. Mellon Foundation Endowed Fund for Innovative Research. We thank Lu Ping Zou for help with sample preparation. The GENESIS model was run on the Cray SVl system at Penn State University's Environment Computing Facility. The manuscript was improved by comments from Thomas Wagner and Paul Pearson and reviews by Paul Wilson and Chris Poulsen. 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(morris@ucsd.edu) BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 LU C3 O LU LU a: o C/3 CO TEMPERATURE (?C) < N o z < z o H LU CD ?B^- CD ^B^"^ cj ^H-H^ --E? 2 12 16 20 24 28 32 405 1 1 1 1 1 1 1 1 1 1 1 1 " V 406 - BenthicT(?C) 16 20 -III 407 408 V ^3 409 ? 4 ; ? Q On 410 411 412 ^ i4 ^ ? " 413 /I-1/1 11CE: 1 1 1 1 1 1 1 1 1 T 1 ??? - ?? ?- : ** - ?^ ? - ? CO - - ^?- D^ ? - 1 1 1 ^Hl'^HPl 1 1 1 7 c I -L < 1^ 05 K 1 m Q U) 1 o. r* 1 < CD yy _i Q: I 1 429 - I I I I I I I I I I ?\7 ^^ 430 h ? V I I I I I I I I I I ? V 0 -1 -2 -3 -4 8^^0 (PDB) -10 12 5^^C (PDB) PLANKTONIC A M. marginata w. baltica BENTHIC O G. volutus D H. moremani Q Praeglob. spp. ? Shallow infaunal ? MIxeci T Hedbergella sp. 0 -^^ bosquensis H. planispira A Deep Infaunal Figure 3. Oxygen and carbon isotope data for DSDP Site 511, cores 47-1 through 49-5. Planktonic foraminifer temperature calculations assume 6^ = ? 1.5%o SMOW. Benthic foraminifer temperature calculations assume 6^ at 1000 m water depth = ? 1.0%o. The horizontal line indicates the position of a subtle lithologic change observed in the archive half of core 511-47-6 at 412.18 mbsf (see text). Other such lithologic changes occur in core 47 but their positions are not noted here. The lithologies are clay stone with rare zeolites (Turonian) and calcareous ooze (Albian). Black fdl represents core recovered. The core number is shown in the core column. 9 - 4 BICE ET AL.: PARADOX OF LATE TURONIAN AT SITE 511 36 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 - ^?^^^^*^?^h^ - 32 S^/ ''K - 28 - VJ^ / ? ^ & _ 24 N\. - T/ V X 1^ _ 20 .^ \ V s \ - \i?? \ \- 16 I \ \ 12 - H H ? 7500 ppm, QFACT=25 ? ? ? ? ? 7500 ppm, QFACT=20 \\ 8 ^^^^ 6500 ppm, QFACT=15 ^l 1 1 1 1 K H ? 4500 ppm, QFACT=12 1 1 1 ?_ 11 4 1 SOS 60S 40S 20S 0 20N 40N 60N SON Latitude O G. volutus W. baltica V Hedbergella sp n H. moremani ^ oyster, James Ross 1. A M. marginata O A- bosquensis Figure 4. Curves show mean January (austral summer) sea surface temperature profiles from four GENESIS version 2.0 AGCM experiments. Symbols show estimated upper ocean temperatures from 8'^0 of approximately coeval late Turonian samples. Isotopic temperature calculations assume the following 6w values: Site 144 foraminifera (5?N lat.) 6w = -0.25%o SMOW; Site 511 foraminifera (60?S lat.) 8w = ? 1.5%o; James Ross Island mollusc (65?S lat.) 8w = ?3.5%o (see text). Symbol shapes for foraminiferal species are the same as those used in Figure 3. All Site 144 estimates are from H. moremani (squares) and Hedbergella sp. (inverted triangles). Site 144 data are from Wilson et al [2002]. The James Ross Island datum is from Ditchfield et al. [1994]. 9 - 7