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T boundary sequences in Denmark and Antarctica
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Earth and Planetary Science Letters 160 (1998) 179–192
Strontium isotope profiles across K=T boundary sequences in Denmark and Antarctica J.M. McArthur a,Ł , M.F. Thirlwall b , M. Engkilde c , W.J. Zinsmeister d , R.J. Howarth a a
Research School of Geological and Geophysical Sciences, University College London and Birkbeck College, Gower Street, London WC1E 6BT, UK b Department of Geology, Royal Holloway and Bedford New College, Egham Hill, Egham, Surrey TW20 0EX, UK c Geological Institute, University of Copenhagen, Øster Voldgade 10, DK-1350 Copenhagen K, Denmark d Department of Earth and Atmospheric Sciences, Purdue University, West Lafayette, IN 47907, USA Received 7 June 1996; accepted 2 April 1998
Abstract Strontium isotope profiles, derived from minimally altered samples, across marine K=T boundary sequences exposed at Kjølby Gaard and Nye Kløv, Denmark, and on Seymour Island, Antarctica, show neither the boundary excursions, nor the boundary spikes, in 87 Sr=86 Sr that have been reported for K=T boundary sequences from elsewhere. Nor do our data conform to modelled predictions of the Sr isotopic response of the oceans to a positive spike in 87 Sr=86 Sr at the terminal Cretaceous. Boundary values for 87 Sr=86 Sr (with 95% confidence intervals) are 0:707828 š 3 in Denmark and 0:707832 š 7 in Antarctica. Our data suggest that 87 Sr=86 Sr stopped increasing and started decreasing at least 90 ka before the K=T boundary. 1998 Elsevier Science B.V. All rights reserved. Keywords: K-T boundary; Sr-87=Sr-86; Antarctica; Denmark
1. Introduction As well as being useful for dating marine sediments (for reviews, see [1–3]), 87 Sr=86 Sr values in marine minerals might be useful for identifying short-term global events [4–12]. In particular, the putative existence of a spike in 87 Sr=86 Sr at the K=T boundary has attracted attention [3–11] because it has been presumed to provide direct evidence of a number of such events, such as: “effects other than those resulting from an impacting meteorite” [4]; the promotion of weathering [5,6] or soil leaching [7] by acidic rain consequent on a meteorite impact; Ł Corresponding
author. E-mail: [email protected]
the result of intense acidic volcanism [9] (a postulate noted by Palmer [11] to be without geological foundation); the effects of a major regression ([10], quoted in Javoy and Courtillot [9]); and the effect of tsunami-driven inputs of continental material to the ocean [8]. Suggestions that the putative K=T spike in 87 Sr=86 Sr reflects the effects of rapid weathering by acidic rain shortly after the K=T boundary events [5–7] draw substance from growing evidence that the Chicxulub structure marks the site of an end-Cretaceous impact [13,14] of an extraterrestrial body into evaporite-bearing rocks [15,16], an event that would have generated large amounts of SO 2 [15– 17] to further acidify rainfall already acidic from
0012-821X/98/$19.00 1998 Elsevier Science B.V. All rights reserved. PII S 0 0 1 2 - 8 2 1 X ( 9 8 ) 0 0 0 5 8 - 2
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NOx dissolution [18]. The magnitude of such a spike might constrain the amount of acidic weathering from acidic gases, whether they result from an impact or volcanic activity [19,20]; according to Martin and Macdougall [6], the end-Cretaceous spike has a magnitude of 28 ð 10 6 , according to Vonhoff and Smit [7] it is 90 ð 10 6 , and according to Meisel et al. [8] it is 130 ð 10 6 (as deduced from their fig. 1). Clearly, small discontinuities in the record of marine 87 Sr=86 Sr might provide important indicators of Earth history, so it is important that the discontinuities be proven to be real before being extensively reported and interpreted. The primary purpose of this paper is to test for the existence of the putative spike=step function in 87 Sr=86 Sr at the K=T boundary, which continues to be evoked in the recent literature [7,21]. A secondary purpose is to establish the 87 Sr=86 Sr value in the K=T boundary interval, in order to further refine the calibration curve of 87 Sr=86 Sr against time for use in strontium isotope stratigraphy [22]. A final purpose is to establish at what time during the latest Maastrichtian=earliest Danian period marine 87 Sr=86 Sr stopped increasing and started to decrease, a matter that might be related to the
Fig. 1. Lithology, biostratigraphy, and sources. Location map inset.
87 Sr=86 Sr
weathering of the Deccan Trap flood basalts [19,20]. To accomplish our aims, we present K=T profiles of 87 Sr=86 Sr for a 15 m sequence of hemipelagic nanno=microfossil chalk at Kjølby Gaard and Nye Kløv, Denmark (Fig. 1), and for a 100 m sequence of fossiliferous, clastic, shallow-marine sediments on Seymour Island, Antarctica (Fig. 2). We assign numerical ages to both sequences to facilitate comparison between them and comparison with data from other authors, despite the uncertainties of such a procedure. We emphasise that our principal finding, that the boundary is not marked by spikes or discontinuities in 87 Sr=86 Sr, is derived from profiles of 87 Sr=86 Sr against stratigraphic level and is independent of the age models adopted for the sequences.
2. Localities: Kjølby Gaard and Nye Kløv 2.1. Lithostratigraphy The Nye Kløv locality is an abandoned chalk quarry in northern Jylland, Denmark that exposes 8 m of upper Maastrichtian and 12 m of lower Danian
profile across the K=T boundary at Kjølby Gaard and Nye Kløv, Denmark. See text for
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Fig. 2. Lithostratigraphy and 87 Sr=86 Sr profile for part of the Lopez de Bertodano Formation, Seymour Island, Antarctica. Inferred relative water depth and positions of glauconite and iridium layers. See text for sources. Location map inset.
chalk (Fig. 1). Sedimentation across the K=T boundary was interrupted at the K=T boundary and at the top of P0 and P1A biozones, but for times that apparently were very short (50 ka; 230 ka, given in Keller et al. [23] compared to the residence time of marine Sr, 2–4 Ma). Maastrichtian sediment is white pelagic chalk containing flint nodules and sparse macrofossils. The Maastrichtian–Danian boundary (K=T boundary) occurs at the base of a 3 cm thick browngrey marly microconglomeratic, smeared, clay. This clay is overlain by 0.5 m of lower Danian grey marly chalk, which is in turn overlain by 8.5 m of white bryozoan chalk with flint nodules and layers. There follows 2.0 m of white bryozoan chalk which grades into white pelagic chalk at the top of the section. The whole section is bioturbated, and flint nodules are common, having formed by early diagenetic silicification of Thallasinoides burrows. Minor tilting of the section, and the microconglomeratic character of the K=T boundary clay, reflect local tectonism related to halokinetic, post-depositional, movement in buried Permian salt domes [24].
2.2. Biostratigraphy Biostratigraphical zonations and lithological data are shown in Fig. 1. The biostratigraphy derives from foraminiferal zonations given by Keller et al. [23] and Bang [25], brachiopod zonations by Surlyk [26] and Johansen [27] and subdivision of the sequence using dinoflagellates [28,29] and coccoliths [30,31]. The uppermost 0.5 m of the Maastrichtian chalk belongs to the Chiropteridium inornatum– Palynodinium grallator dinoflagellate zone, which is restricted to the central part of the Danish basin [29].
3. Localities: Seymour Island 3.1. Lithostratigraphy The sequence on Seymour Island comprises the upper parts of the Lopez de Bertodano Formation, and the overlying Sobral Formation (Fig. 2). The for-
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mer comprises some 1190 m of grey or tan, friable, sandy to muddy siltstones, which have been divided into ten units [32], the upper four (7–10) of which are interpreted to be middle and outer shelf facies [32]. The overlying Sobral Formation was deposited by a prograding deltaic system [32], and comprises 255 m of maroon, laminated, sediments that grade upwards from basal silts into cleaner sandstones. The sequence becomes increasingly glauconitic and more cross-bedded towards the top. The K=T boundary sequence in the Lopez de Bertodano Formation appears to be a marine record that is as nearly continuous across the boundary as is likely to be found. It is greatly expanded relative to abyssal sequences [33–35]. Our numeric calibration of the sequence (see later) suggests that it accumulated at a rate of about 27 cm ka 1 (compacted), a rate up to 4 times faster than that in Denmark and between 10 and 80 times faster than for the sequences examined by Martin and Macdougall [6]. 3.2. Biostratigraphy Details of faunal distributions for Seymour Island have been presented elsewhere [33,34,36]. The upper units (8–10) of the Lopez de Bertodano Formation contain abundant bivalves, gastropods and ammonites [32,33]. The high sedimentation rate has preserved a detailed record of Late Cretaceous and early Danian fossil evolution and extinction [33]. Owing to the expanded nature of the sequence, faunal indicators used to define the K=T boundary, such as planktonic foraminifera, ammonites and dinocysts, that are commonly coincident in more condensed sequences, disappear at different levels within a 30 m boundary interval that spans the contact between units 9 and 10 of the Lopez de Bertodano Formation (Fig. 2). No single extinction horizon is particularly distinguished [33]; the boundary position depends upon the biological marker used to define it. Consequently, the boundary has more recently been placed at the level of an iridium-enriched layer within a glauconitic bed about 9 m above the contact of units 9 and 10 of the Lopez de Bertodano Formation and we adopt this boundary definition here [33,35,37]. This placement is some 0.5 to 1 m below the K=T boundary as defined by dinocysts and about 2 m above the highest in-place
ammonite [33,37]. The overlying Sobral Formation is assigned a Lower Paleocene (Danian) age based on dinoflagellates [38] and foraminifera [36].
4. Samples Samples from Denmark comprise 28 samples of nannofossil chalk collected at Nye Kløv, a section 400 m east of a quarry at Kjølby Gaard on the north side of Limfjorden in northwestern Jylland. In addition, eleven nannofossil chalk samples were collected from the section at Kjølby Gaard (Table 1; Fig. 1). Stratigraphic positions of samples are reported as metres above or below the K=T boundary. Samples from Seymour Island comprised 23 bivalves, gastropods, serpulids, echinoids, and a belemnite collected from units 8, 9 and 10 of the Lopez de Bertodano Formation and from the overlying Sobral Formation [33]. Details of samples, including their stratigraphic levels, are given in Table 1 and Fig. 2. Stratigraphic levels are given in metres above or below the K=T boundary.
5. Analytical methods Analytical methods used were those in [39–41] for bulk chalk and in [42] for macrofossils and are only summarised here. Apparently altered portions of macrofossils were removed using diamond cutting tools; the remaining sample was broken into mm-sized fragments. These were cleaned by brief immersion in 1.2 M hydrochloric acid solution, then ultrapure water, and finally dried in a clean environment. Fragments for analysis were then picked under the binocular microscope. Chalk samples were prepared by disaggregation in water using ultrasound, followed by dissolution and disposal of 30% of the sample in dilute acetic acid. After washing and centrifugation, the acid-leached portion of the sample was analysed for 87 Sr=86 Sr. This procedure removes much of the contaminant Sr, and Sr in diagenetic overgrowths, leaving an 87 Sr=86 Sr value representative of the original marine value [39–41]. For eighteen samples, comparison was made of 87 Sr=86 Sr values after one pre-leach with those measured after two pre-leaches. The mean difference was 6 ð 10 6 ,
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a result that is considered to show that no significant imprint of diagenesis remained after one pre-leach. The samples were screened for alteration using a combination of SEM, XRD, and chemical analysis. Chalk samples were examined by SEM in order to assess the amount of diagenetic cement, spar, and overgrowths on sediment components. The resulting data are given in Table 1, whilst some of the ultrastructural information obtained using the SEM is shown in Fig. 3. Chemical analysis for Ca, Mg, Fe, and Mn was done by standard methods using ICP–AES (Ca, Sr, Mg, Fe, Mn) and furnace AAS (Rb) on samples dissolved in 1.8 M acetic acid. The precision of the analysis was better than š10%. Ultrastructure was examined under gold coatings using a Zeiss DSM 940 SEM. To determine mineralogy, samples were analysed using a Phillips PW 1710 Diffractometer. Instrumental conditions were: CuKα radiation generated at 40 kV and 30 mA; scanning through 24– 32º 2Q at 0.5º=min, i.e. for the major diffraction peaks of aragonite, calcite, and dolomite. With these conditions the detection limit of our apparatus, determined by standard additions with pure phases, was about 0.2% calcite in aragonite and 0.5% dolomite in calcite. For 87 Sr=86 Sr analysis, cleaned macrofossil fragments were dissolved in 2.5 M hydrochloric or 1.8 M acetic acid. These preparations gave indistinguishable data. Pre-leached chalks were disaggregated with ultrasound in 2 ml of pure water and 30% of the sample was then dissolved by addition of 400 µl of 1.8 M acetic acid. After reaction, the solution was isolated for analysis by double centrifugation. Remaining undissolved sample was discarded. Values of 87 Sr=86 Sr were determined with a VG-354 fivecollector mass spectrometer using the multi-dynamic routines that include corrections for isobaric interference from 87 Rb [43]. Data have been normalised to a value of 0.1194 for 86 Sr=88 Sr. The data were collected between April, 1994, and September, 1997. During data collection, the measured value for NIST 987 was within 0.00002 of the value 0.710248. Data reported in Table 1 have been adjusted to a value of 0.710248 for NIST 987 which, in our laboratory corresponds to a value of 0.709175 for Modern Seawater Sr (MSS). Based upon replicated sample analysis the precision of our measurements (2 s.e.)
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was š15 ð 10 6 for single determinations, š11 for duplicates and š9 for triplicates. Total blanks were <2 ng of Sr. Sample contained >5 mg of Sr. Concentrations of Rb were found to be too low to require corrections for radiogenic Sr. Data from the literature that are reported here are adjusted to a value of 0.710248 for NIST 987.
6. Results and discussion 6.1. Sample integrity and diagenesis As all our samples have been buried at some time, they have been affected by diagenesis. We believe, for the reasons given below, that the alteration to 87 Sr=86 Sr is too small to affect our conclusions. Perhaps the most telling reason for our belief is that 87 Sr=86 Sr values in the K=T boundary (with values and 95% confidence intervals obtained by the non-parametric regression method given in [22]) are 0:707828 š 3 for Denmark and 0:707832 š 7 for Seymour Island, with the trend of the data from each locality being concordant (Figs. 1 and 2). Such good agreement, for different materials from opposite ends of the Earth, is unlikely to be fortuitous. It occurs only because each sequence is recording the original 87 Sr=86 Sr in K=T seawater. Other evidence supports this conclusion. Most importantly, macrofossils from Seymour Island retain their ultrastructural integrity, which suggests that they are effectively unaltered. Additionally, we have not analysed bulk chalk from Denmark; we have analysed about half of that part of a bulk sample that remains after an acid pre-leach has removed 30% of the sample (overgrowths, exchangeable Sr, acid-soluble clays). We expand on these points below and in the methods section. Firstly, chalk from northwestern Germany and eastern England yield excellent records of 87 Sr=86 Sr isotope evolution for the Late Cretaceous [39– 41,44], unquestionably record original 87 Sr=86 Sr values (when analysed by the method used here i.e. pre-leached) and are similar in appearance, chemical composition (Table 1), and diagenetic state (Fig. 3) to, and come from the same basin as, our Danish chalks. The effectiveness of our pre-leach method is clear from 87 Sr=86 Sr analysis of horizons of chalk at Hemmoor, northwestern Germany [44] in which
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Table 1 Chemical and isotopic data for K=T boundary sequences, Kjølby Gaard and Nye Klov, Denmark, and Seymour Island, Antarctica. Samples are macrofollis (SI) and acid-leached nannofossil chalks (KG, NK)
J.M. McArthur et al. / Earth and Planetary Science Letters 160 (1998) 179–192 Table 1 (continued).
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Fig. 3. SEM photomicrographs showing preservation of samples. (a) Chalk from Lagerdorf (sample 88 of [41]) that preserves a good record of original 87 Sr=86 Sr. (b) Chalk (KG12) from Kjølby Gaard. In (a) and (b) overgrowth is present in both but spar is rare. (c) and (d) Pycnodonte samples (c, 1175; d, 1163) from Seymour Island, showing unaltered compact layering in calcite. (e) Good preservation of Pseudotextularia deformis (long axis 200 µm) from 1 m below the K=T boundary, Kjølby Gaard. (f) Detail of aperture in (e) showing only minor nucleation of calcite on the specimen’s surface.
localised diagenetic cementation by carbonate has accompanied flint formation. Compared to nearby uncemented chalks, cemented horizons have 40%
lower Sr=Ca, but 87 Sr=86 Sr values (of pre-leached samples) that differ by less than 10 ð 10 6 (fig. 4 of [44]). Furthermore, samples of pre-leached nanno-
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fossil chalk (KG11) and macrofossil debris (KG12) from the same stratigraphic level in Kjølby Gaard have 87 Sr=86 Sr values that are analytically indistinguishable (Table 1). Finally, for twelve samples, we compared 87 Sr=86 Sr using pre-leaches that dissolved and discarded either 30% or 60% of the sample prior to analysis; the mean difference was 6 ð 10 6 , which suggests that our data are not significantly affected by diagenetic artefacts. The Seymour Island samples (all macrofossils) are monomineralic and have concentrations of Sr, Fe and Mn (Table 1) in the range shown by numerous authors ([3,45], and refs. therein) to be representative of well-preserved macrofossils. Ultrastructural detail is well preserved (Fig. 3). Replicate analyses on separate subsamples of macrofossil material show excellent agreement; this would not be expected were preservation poor. Finally, our macrofossils comprise a variety of faunal types, some retaining an original aragonitic mineralogy, whilst others retain their original calcite. Were these fossils altered, it would be unlikely that such different organisms would all alter to yield such similar isotopic ratios. Finally, samples considered poorly preserved by virtue of their mixed mineralogy, high Fe and Mn content, or anomalously high Sr content (Table 1) generally have an 87 Sr=86 Sr value much lower than that expected for the interval: in these sediments, alteration of macrofossil carbonate in our samples lowers 87 Sr=86 Sr rather than increasing it, as is usual. 6.2. Trends in 87 Sr=86 Sr When plotted against stratigraphic level (Figs. 1 and 2), 87 Sr=86 Sr changes by less than 20 ð 10 6 from the base (0.70784) to the top (0.70782) of the sequences. Statistical regression [22] of the data defines boundary values (with 95% CIs) as 0:707828 š 3 in Denmark and 0:707832 š 7 in Antarctica. Neither sequence shows positive boundary discontinuities in 87 Sr=86 Sr, as has been reported elsewhere for K=T sequences [4–8]. In both, however, there is a suggestion in the trends that values decline across the boundary and reach a minimum of 0.70782 just after it (about 0.5–1 m in Denmark; 20–40 m in Antarctica), thereafter increasing up the sequence through a broad maximum of 0.70783 before resuming the downward trend seen in the lower
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parts of the sequences. This trend is best seen in the Danish data, which are presumed to have a more extended post-boundary stratigraphic range than do the data from Antarctica. These data do not confirm the existence of a positive excursion in 87 Sr=86 Sr at the K=T boundary. A geological catastrophe, such as a meteorite impact, sudden climate change, sudden inundation of arid salt-laden continental areas, sudden glacial events, periods of acidic rain, etc., may well cause a sudden change in the 87 Sr=86 Sr of seawater, but, as correctly noted by Martin and Macdougall [6] and Nelson et al. [46], such excursions cannot suddenly disappear from seawater (and so the rock record) but must decay to the underlying trend in 87 Sr=86 Sr with a time constant that is the reciprocal of the residence time of marine Sr. Martin and Macdougall [6] point out that, with a residence time for Sr of 4 Ma, their putative spike should have decayed by only 20–25% of its value after 1 Ma, but that it had actually returned to pre-K=T values by that time, a puzzle they left unresolved (as they did the origin of the spike in 87 Sr=86 Sr of 45 ð 10 6 shown by their fig. 4 to occur about 2 Ma after the K=T boundary). It follows that any ‘spike’ with a duration of much less than 0.5 Ma must be an artefact of diagenesis or analysis or, if real, must decay exponentially to the background trend with a time constant that can reveal the true residence time of Sr. In Fig. 4a, we plot 87 Sr=86 Sr data against numerical age for our data and that of Martin and Macdougall [6], both superimposed on the modelled decay curve from a positive spike size of 28 ð 106 , the size reported by Martin and Macdougall [6]. The figure incorporates the underlying Maastrichtian increase of 19 ð 10 6 Ma 1 [47] and a Danian decrease of 15 ð 10 6 Ma 1 [22]. The residence time for Sr of 2 Ma was chosen to represent the shortest (and so fastest decay) thought reasonable, and that of 0.5 Ma to show the insensitivity of our conclusions to extreme choices. We have assigned numerical ages to the Seymour Island samples using an 87 Sr=86 Sr boundary value of 0.707832, a K=T boundary date of 65.0 Ma [51], and the fact that 1400 m of Maastrichtian sediments were deposited on Seymour Island and Snow Hill Island (A. Crame, pers. commun., 1997) during its 6.3 Ma [52] duration. We assume that sedimentation rate was constant
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Fig. 4. (a) Comparison of data with modelled 87 Sr=86 Sr decay curves. The model assumes an instantaneous spike of C28 ð 106 in 87 Sr=86 Sr at the K=T boundary [6], residence time for Sr of either 2 or 0.5 Ma and exponential decay of the 87 Sr=86 Sr spike according to the formula [87 Sr=86 Sr] D [87 Sr=86 Sr]o exp. t=RT/, where t D elapsed time, RT D oceanic residence time of Sr, and [87 Sr=86 Sr]o is the spike magnitude. Decay is superimposed on background trends for the late Maastrichtian of C0.000019 Ma 1 [47] and 0.000015 Ma 1 for the Danian [22]. Filled triangles are Seymour Island; filled circles are Denmark; open symbols from Martin and Macdougall [6], numerical ages for Sites 528 and 577 have been adjusted to include boundary hiatuses of 60 ka not recognised by them but reported elsewhere [48]. Data normalised to 0.710248 for SRM (NIST) 987. (b) Compilation of 87 Sr=86 Sr data as a function of numerical age for K=T samples reported here (filled circles) and those from Hess et al. [4] (hourglasses), Martin and Macdougall [6] (open circles), McArthur et al. [41] (open squares), McLaughlin et al. [44] (crosses), Sugarman et al. [47] (open triangles), Dennison et al. [49] (crossed squares), DePaolo and Ingram [50] (inverted triangles). All numeric ages have been adjusted to a common timescale where the K=T boundary is 65.0 Ma. Data of Hess et al. [4] adjusted using a value for their MSS of 0.709185, computed from their table 2 after exclusion of an obvious flyer, rather than their incorrectly computed value of 0.709198. Data normalised to 0.710248 for SRM (NIST) 987, except for Hess et al. [4], which is normalised to 0.709175 for MSS.
through the 1400 m of the Maastrichtian sequence. To the Danian part of the Danish sequence, we use the age model for Nye Kløv given by Keller et al. [23]. To the Maastrichtian part of the sequence we use a value of 4 cm ka 1 , one midway between the mean rate of 9 cm ka 1 for the entire Maastrichtian at its thickest in the Danish Basin [Kaminski 49], but viewed as too high by Keller et al. [23], and the values for several abyssal sites during K=T times of 0.3 to 2.5 cm 1 ka 1 [6]. Fig. 4a shows that our data fail to conform to the modelled decay curve, but do fit the underlying Maastrichtian and Danian trends. The data of Martin and Macdougall [6] scatter more widely than do ours; this scatter compromises a loose correspondence of Martin and Macdougall’s Danian data [6] to the modelled decay curve. Inter-laboratory bias of some 17 ð 10 6 in excess of that which can be accounted for by normalisation to standards has been noted to occur between laboratories [3,53]; whether this affects our comparison is not known. Data for K=T boundary strata from elsewhere are sparse, but are generally concordant with ours (Fig. 4b). Data for a Maastrichtian interval at Bidart, southwestern France [46] showing a boundary value of 0.707842 at the K=T boundary is in question after the discovery that foraminifera from this sequence have shells that luminesce under CL and contain high concentrations of iron (H. Vonhoff, pers. commun., 1996), facts that suggest to us that they have been altered despite preserving good external morphologies. A boundary sequence at Braggs, Alabama [54] shows a broad maximum, but no spike, through about 4 m of section centred on the boundary (boundary value about 0.70784 derived from fig. 1 of [54], corrected for interlaboratory bias [55]. The Sr-isotope regression of Sugarman et al. [47] assumes 87 Sr=86 Sr increases linearly to the boundary; their regression gives a value at 65.1 Ma of 0.707845 (at our postulated Maastrichtian maximum) but shows no discontinuities in the few data that define their trend prior to it (Fig. 4b for data points). Extrapolated to 65.0 Ma using a post-maximum decrease in 87 Sr=86 Sr of 15 ð 10 6 , their data predict a value of 0.707844 for the K=T boundary, a similar result to that given by these authors for a boundary sample from El Kef (0.707847 at 65.1 Ma).
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6.3. An explanation for spikes We believe that spikes and step functions in existing 87 Sr=86 Sr compilations for K=T boundary intervals result from analytical and diagenetic effects. The original report of a boundary spike, of 100 ð 10 6 , was made by Hess et al. [4] on the basis of two anomalous values (Fig. 4b). Neither measurement was duplicated; without such duplication neither data can be relied upon. The samples were 29-3-5-7 at Site 356, which these authors omitted to categorise for preservational quality (their table 2), and sample 12-5-19-21 from Site 577, for which there was no evidence of recrystallisation. Martin and Macdougall [6] report a maximum as occurring about 60 ka after the boundary, but they clearly state that preservation of samples was different on either side of the boundary at all the localities they examined. For example, for Site 356, their most densely sampled locality and the only one with a convincing discontinuity across the boundary, they state that “Maastrichtian samples frequently contain calcite overgrowths on their interior walls, while Danian foraminifera exhibit excellent preservation.” Diagenesis is expected to have affected boundary sediments at their sites as all have a hiatus at the boundary of at least 50 ka 1 [6,48]. The discontinuity in 87 Sr=86 Sr at Site 356 may result from diagenetic lowering of 87 Sr=86 Sr below the boundary. Meisel et al. [8] report a spike of 130 ð 10 6 across the K=T boundary at Sumbar, Turkmenistan. To isolate Sr from their samples, on the boundary clay itself, they were attacked with 5 M acetic acid without being pre-leached to remove contaminant Sr. Contrary to the statement of Meisel et al. that “leaching with 5 M acetic acid leaves clay minerals unaltered”, it leaves them leached of some of their radiogenic Sr; this treatment minimises, but does not prevent, dissolution of contaminant Sr from clay minerals by the hydrogen ion [3]. It is not surprising that their 87 Sr=86 Sr values appear to be lithology dependent; for those lithologies that might reasonably be expected to give the least contaminated ratios, because their Sr comes mostly from carbonate (marlstone S9, 13 to 16 cm above the boundary; marlstone S43, 0 to 1 cm below it), the mean ratio is 0:707834 š 10, a value close to our boundary values.
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The spike in 87 Sr=86 Sr reported by Meisel et al. is an analytical artefact. Attempts to use acids to isolate Sr from carbonate in clay-rich sediments are unlikely to succeed. Finally, Vonhoff and Smit [7] also report a boundary spike in 87 Sr=86 Sr around the K=T boundary. Most of their samples are rich in iron and manganese and luminesce brightly under CL, facts which compromise the reliability of their data by suggesting that their samples are altered. Those that are not apparently altered (two visually well preserved foraminifera from the E-clay, Guelhemmerburg Quarry) contain enough diagenetic carbonate (up to 1%) to yield the spike they report, even on the most conservative estimates of the Sr abundance and isotopic ratio in that diagenetic calcite, given that the foraminifera have resided for 65 Ma in a clay-rich (i.e. radiogenic) unit with a bulk 87 Sr=86 Sr of 0.75 (H. Vonhoff, pers. commun., 1998). The 87 Sr=86 Sr of Sr in seawater stopped increasing and started decreasing close to the K=T boundary [4,6,2,3,22]. Such a reversal requires a major change in the balance of Sr supplied to the ocean by mid-ocean ridges and rivers. We agree with the speculation of Vonhoff and Smit [7] that this decrease may have coincided with Deccan Trap volcanism and its consequences, but note that the downturn in 87 Sr=86 Sr at this time marked the end of a period of increase that lasted 25 Ma and the beginning of a period of decrease that lasted for 13 Ma [22]. The underlying reason for the change must therefore relate to mantle adjustments to the outpouring of the Deccan basalts, rather than the consequences of their surface weathering. Unfortunately, the number of variables, and the lack of quantification of their magnitudes, needed to construct models that address why such changes occur in the marine 87 Sr=86 Sr record preclude a useful discussion of this issue, a conclusion broadly in agreement with that of a few others [56,57].
7. Conclusions The mean 87 Sr=86 Sr at the K=T boundary, as determined from our combined data using non-parametric regression [22], is estimated as 0:707830 š 4 (95% CI), which is close to values found by others
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for this boundary. Our data do not show a spike or a step function in 87 Sr=86 Sr at the K=T boundary. The K=T boundary discontinuities in 87 Sr=86 Sr that have been reported elsewhere are artefacts of diagenesis or analytical methodology, or both. Marine 87 Sr=86 Sr peaked at or before 65.1 Ma and decreased thereafter into the Danian.
Acknowledgements The Radiogenic Isotope Laboratory at RHBNC is supported, in part, by the University of London as an ´ .M. McLaughlin intercollegiate facility. We thank O for assisting with the isotopic measurements, A. Osborn for assistance with sample preparation and elemental analysis, T. Styles and C. Stuart for help with preparation of figures and J. Davy for the SEM work. Samples from Kjølby Gaard were provided by A.S. Gale. We thank Finn Surlyk, Gerta Keller, Donald DePaolo, and two anonymous reviewers for constructive critiques of the script. [FA]
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J.M. McArthur et al. / Earth and Planetary Science Letters 160 (1998) 179–192 [26] F. Surlyk, The Maastrichtian Stage in NW Europe, and its brachiopod zonation, Bull. Geol. Soc. Denm. 33 (1984) 1– 2, 217–224. [27] M.B. Johansen, Brachiopods from the Maastrichtian– Danian boundary sequence at Nye-Kløv, Jylland Denmark, Fossils Strata 20 (1987) 1–99. [28] J.M. Hansen, Dinoflagellate stratigraphy and echinoid distribution in Upper Maastrichtian and Danian deposits from Denmark, Bull. Geol. Soc. Denm. 26 (1977) 1–26. [29] J.M. Hansen, Dinoflagellate zonation around the boundary, in: T. Birkelund, R.G. Bromley (Eds.), Cretaceous–Tertiary Boundary Events, 1, The Maastrichtian and Danian of Denmark, Univ. Copenhagen, 1979, pp. 136–141. [30] K. Perch-Nielsen, Calcareous nannofossil zonation at the Cretaceous–Tertiary boundary in Denmark, in: T. Birkelund, R.G. Bromley (Eds.), Symposium on Cretaceous–Tertiary Boundary Events, 1, The Maastrichtian and Danian of Denmark, Univ. Copenhagen, 1979, pp. 115– 135. [31] K. Perch-Nielsen, Calcareous nannofossils in Cretaceous– Tertiary boundary sections in Denmark, in: W.K. Christensen, T. Birkelund (Eds.), Symposium on Cretaceous– Tertiary Boundary Events, 2, Univ. Copenhagen, 1979, pp. 120–126. [32] C.E. Macellari, Stratigraphy, sedimentology, and paleoecology of Upper Cretaceous=Paleocene shelf-deltaic sediments of Seymour Island, Geol. Soc. Am., Mem. 169 (1988) 25– 53. [33] W.J. Zinsmeister, R.M. Feldmann, M.O. Woodburne, D.H. Elliot, Latest Cretaceous=earliest Tertiary transition on Seymour Island, Antarctica, J. Paleontol. 63 (1989) 731–738. [34] B.T. Huber, Late Cretaceous Foraminiferal Biostratigraphy, Paleoecology, and Paleobiogeography of the James Ross Island Region, Antarctic Peninsula, unpubl. M.Sc. thesis, The Ohio State Univ., Columbus, 1984, 246 pp. [35] D.H. Elliot, R.A. Askin, F.T. Kyte, W.J. Zinsmeister, Iridium and dinocysts at the Cretaceous–Tertiary boundary on Seymour Island, Antarctica: implications for the K–T event, Geology 22 (1994) 675–678. [36] B.T. Huber, Upper Campanian–Paleocene foraminifera from the James Ross Island region, Antarctic Peninsula, Geol. Soc. Am. Mem. 169 (1988) 163–252. [37] C.R. Marshall, Distinguishing between sudden and gradual extinctions in the fossil record: Predicting the position of the Cretaceous–Tertiary iridium anomaly using the ammonite fossil record on Seymour Island, Antarctica, Geology 23 (1995) 731–734. [38] D.M. Harwood, Upper Cretaceous and lower Paleocene diatom and silicoflagellate biostratigraphy of Seymour Island, eastern Antarctic Peninsula, Geol. Soc. Am. Mem. 16 (1988) 55–130. [39] J.M. McArthur, A.S. Gale, W.J. Kennedy, M. Chen, M.F. Thirlwall, Strontium-isotope stratigraphy in the Late Cretaceous: international correlation of the Campanian=Maastrichtian boundary, Terra Nova 4 (1992) 332–345. [40] J.M. McArthur, M.F. Thirlwall, A.S. Gale, W.J. Kennedy, J.A. Burnett, D. Mattey, A.R. Lord, Strontium isotope
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Strontium isotope profiles across K=T boundary sequences in Denmark and Antarctica J.M. McArthur a,Ł , M.F. Thirlwall b , M. Engkilde c , W.J. Zinsmeister d , R.J. Howarth a a
Research School of Geological and Geophysical Sciences, University College London and Birkbeck College, Gower Street, London WC1E 6BT, UK b Department of Geology, Royal Holloway and Bedford New College, Egham Hill, Egham, Surrey TW20 0EX, UK c Geological Institute, University of Copenhagen, Øster Voldgade 10, DK-1350 Copenhagen K, Denmark d Department of Earth and Atmospheric Sciences, Purdue University, West Lafayette, IN 47907, USA Received 7 June 1996; accepted 2 April 1998
Abstract Strontium isotope profiles, derived from minimally altered samples, across marine K=T boundary sequences exposed at Kjølby Gaard and Nye Kløv, Denmark, and on Seymour Island, Antarctica, show neither the boundary excursions, nor the boundary spikes, in 87 Sr=86 Sr that have been reported for K=T boundary sequences from elsewhere. Nor do our data conform to modelled predictions of the Sr isotopic response of the oceans to a positive spike in 87 Sr=86 Sr at the terminal Cretaceous. Boundary values for 87 Sr=86 Sr (with 95% confidence intervals) are 0:707828 š 3 in Denmark and 0:707832 š 7 in Antarctica. Our data suggest that 87 Sr=86 Sr stopped increasing and started decreasing at least 90 ka before the K=T boundary. 1998 Elsevier Science B.V. All rights reserved. Keywords: K-T boundary; Sr-87=Sr-86; Antarctica; Denmark
1. Introduction As well as being useful for dating marine sediments (for reviews, see [1–3]), 87 Sr=86 Sr values in marine minerals might be useful for identifying short-term global events [4–12]. In particular, the putative existence of a spike in 87 Sr=86 Sr at the K=T boundary has attracted attention [3–11] because it has been presumed to provide direct evidence of a number of such events, such as: “effects other than those resulting from an impacting meteorite” [4]; the promotion of weathering [5,6] or soil leaching [7] by acidic rain consequent on a meteorite impact; Ł Corresponding
author. E-mail: [email protected]
the result of intense acidic volcanism [9] (a postulate noted by Palmer [11] to be without geological foundation); the effects of a major regression ([10], quoted in Javoy and Courtillot [9]); and the effect of tsunami-driven inputs of continental material to the ocean [8]. Suggestions that the putative K=T spike in 87 Sr=86 Sr reflects the effects of rapid weathering by acidic rain shortly after the K=T boundary events [5–7] draw substance from growing evidence that the Chicxulub structure marks the site of an end-Cretaceous impact [13,14] of an extraterrestrial body into evaporite-bearing rocks [15,16], an event that would have generated large amounts of SO 2 [15– 17] to further acidify rainfall already acidic from
0012-821X/98/$19.00 1998 Elsevier Science B.V. All rights reserved. PII S 0 0 1 2 - 8 2 1 X ( 9 8 ) 0 0 0 5 8 - 2
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NOx dissolution [18]. The magnitude of such a spike might constrain the amount of acidic weathering from acidic gases, whether they result from an impact or volcanic activity [19,20]; according to Martin and Macdougall [6], the end-Cretaceous spike has a magnitude of 28 ð 10 6 , according to Vonhoff and Smit [7] it is 90 ð 10 6 , and according to Meisel et al. [8] it is 130 ð 10 6 (as deduced from their fig. 1). Clearly, small discontinuities in the record of marine 87 Sr=86 Sr might provide important indicators of Earth history, so it is important that the discontinuities be proven to be real before being extensively reported and interpreted. The primary purpose of this paper is to test for the existence of the putative spike=step function in 87 Sr=86 Sr at the K=T boundary, which continues to be evoked in the recent literature [7,21]. A secondary purpose is to establish the 87 Sr=86 Sr value in the K=T boundary interval, in order to further refine the calibration curve of 87 Sr=86 Sr against time for use in strontium isotope stratigraphy [22]. A final purpose is to establish at what time during the latest Maastrichtian=earliest Danian period marine 87 Sr=86 Sr stopped increasing and started to decrease, a matter that might be related to the
Fig. 1. Lithology, biostratigraphy, and sources. Location map inset.
87 Sr=86 Sr
weathering of the Deccan Trap flood basalts [19,20]. To accomplish our aims, we present K=T profiles of 87 Sr=86 Sr for a 15 m sequence of hemipelagic nanno=microfossil chalk at Kjølby Gaard and Nye Kløv, Denmark (Fig. 1), and for a 100 m sequence of fossiliferous, clastic, shallow-marine sediments on Seymour Island, Antarctica (Fig. 2). We assign numerical ages to both sequences to facilitate comparison between them and comparison with data from other authors, despite the uncertainties of such a procedure. We emphasise that our principal finding, that the boundary is not marked by spikes or discontinuities in 87 Sr=86 Sr, is derived from profiles of 87 Sr=86 Sr against stratigraphic level and is independent of the age models adopted for the sequences.
2. Localities: Kjølby Gaard and Nye Kløv 2.1. Lithostratigraphy The Nye Kløv locality is an abandoned chalk quarry in northern Jylland, Denmark that exposes 8 m of upper Maastrichtian and 12 m of lower Danian
profile across the K=T boundary at Kjølby Gaard and Nye Kløv, Denmark. See text for
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Fig. 2. Lithostratigraphy and 87 Sr=86 Sr profile for part of the Lopez de Bertodano Formation, Seymour Island, Antarctica. Inferred relative water depth and positions of glauconite and iridium layers. See text for sources. Location map inset.
chalk (Fig. 1). Sedimentation across the K=T boundary was interrupted at the K=T boundary and at the top of P0 and P1A biozones, but for times that apparently were very short (50 ka; 230 ka, given in Keller et al. [23] compared to the residence time of marine Sr, 2–4 Ma). Maastrichtian sediment is white pelagic chalk containing flint nodules and sparse macrofossils. The Maastrichtian–Danian boundary (K=T boundary) occurs at the base of a 3 cm thick browngrey marly microconglomeratic, smeared, clay. This clay is overlain by 0.5 m of lower Danian grey marly chalk, which is in turn overlain by 8.5 m of white bryozoan chalk with flint nodules and layers. There follows 2.0 m of white bryozoan chalk which grades into white pelagic chalk at the top of the section. The whole section is bioturbated, and flint nodules are common, having formed by early diagenetic silicification of Thallasinoides burrows. Minor tilting of the section, and the microconglomeratic character of the K=T boundary clay, reflect local tectonism related to halokinetic, post-depositional, movement in buried Permian salt domes [24].
2.2. Biostratigraphy Biostratigraphical zonations and lithological data are shown in Fig. 1. The biostratigraphy derives from foraminiferal zonations given by Keller et al. [23] and Bang [25], brachiopod zonations by Surlyk [26] and Johansen [27] and subdivision of the sequence using dinoflagellates [28,29] and coccoliths [30,31]. The uppermost 0.5 m of the Maastrichtian chalk belongs to the Chiropteridium inornatum– Palynodinium grallator dinoflagellate zone, which is restricted to the central part of the Danish basin [29].
3. Localities: Seymour Island 3.1. Lithostratigraphy The sequence on Seymour Island comprises the upper parts of the Lopez de Bertodano Formation, and the overlying Sobral Formation (Fig. 2). The for-
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mer comprises some 1190 m of grey or tan, friable, sandy to muddy siltstones, which have been divided into ten units [32], the upper four (7–10) of which are interpreted to be middle and outer shelf facies [32]. The overlying Sobral Formation was deposited by a prograding deltaic system [32], and comprises 255 m of maroon, laminated, sediments that grade upwards from basal silts into cleaner sandstones. The sequence becomes increasingly glauconitic and more cross-bedded towards the top. The K=T boundary sequence in the Lopez de Bertodano Formation appears to be a marine record that is as nearly continuous across the boundary as is likely to be found. It is greatly expanded relative to abyssal sequences [33–35]. Our numeric calibration of the sequence (see later) suggests that it accumulated at a rate of about 27 cm ka 1 (compacted), a rate up to 4 times faster than that in Denmark and between 10 and 80 times faster than for the sequences examined by Martin and Macdougall [6]. 3.2. Biostratigraphy Details of faunal distributions for Seymour Island have been presented elsewhere [33,34,36]. The upper units (8–10) of the Lopez de Bertodano Formation contain abundant bivalves, gastropods and ammonites [32,33]. The high sedimentation rate has preserved a detailed record of Late Cretaceous and early Danian fossil evolution and extinction [33]. Owing to the expanded nature of the sequence, faunal indicators used to define the K=T boundary, such as planktonic foraminifera, ammonites and dinocysts, that are commonly coincident in more condensed sequences, disappear at different levels within a 30 m boundary interval that spans the contact between units 9 and 10 of the Lopez de Bertodano Formation (Fig. 2). No single extinction horizon is particularly distinguished [33]; the boundary position depends upon the biological marker used to define it. Consequently, the boundary has more recently been placed at the level of an iridium-enriched layer within a glauconitic bed about 9 m above the contact of units 9 and 10 of the Lopez de Bertodano Formation and we adopt this boundary definition here [33,35,37]. This placement is some 0.5 to 1 m below the K=T boundary as defined by dinocysts and about 2 m above the highest in-place
ammonite [33,37]. The overlying Sobral Formation is assigned a Lower Paleocene (Danian) age based on dinoflagellates [38] and foraminifera [36].
4. Samples Samples from Denmark comprise 28 samples of nannofossil chalk collected at Nye Kløv, a section 400 m east of a quarry at Kjølby Gaard on the north side of Limfjorden in northwestern Jylland. In addition, eleven nannofossil chalk samples were collected from the section at Kjølby Gaard (Table 1; Fig. 1). Stratigraphic positions of samples are reported as metres above or below the K=T boundary. Samples from Seymour Island comprised 23 bivalves, gastropods, serpulids, echinoids, and a belemnite collected from units 8, 9 and 10 of the Lopez de Bertodano Formation and from the overlying Sobral Formation [33]. Details of samples, including their stratigraphic levels, are given in Table 1 and Fig. 2. Stratigraphic levels are given in metres above or below the K=T boundary.
5. Analytical methods Analytical methods used were those in [39–41] for bulk chalk and in [42] for macrofossils and are only summarised here. Apparently altered portions of macrofossils were removed using diamond cutting tools; the remaining sample was broken into mm-sized fragments. These were cleaned by brief immersion in 1.2 M hydrochloric acid solution, then ultrapure water, and finally dried in a clean environment. Fragments for analysis were then picked under the binocular microscope. Chalk samples were prepared by disaggregation in water using ultrasound, followed by dissolution and disposal of 30% of the sample in dilute acetic acid. After washing and centrifugation, the acid-leached portion of the sample was analysed for 87 Sr=86 Sr. This procedure removes much of the contaminant Sr, and Sr in diagenetic overgrowths, leaving an 87 Sr=86 Sr value representative of the original marine value [39–41]. For eighteen samples, comparison was made of 87 Sr=86 Sr values after one pre-leach with those measured after two pre-leaches. The mean difference was 6 ð 10 6 ,
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a result that is considered to show that no significant imprint of diagenesis remained after one pre-leach. The samples were screened for alteration using a combination of SEM, XRD, and chemical analysis. Chalk samples were examined by SEM in order to assess the amount of diagenetic cement, spar, and overgrowths on sediment components. The resulting data are given in Table 1, whilst some of the ultrastructural information obtained using the SEM is shown in Fig. 3. Chemical analysis for Ca, Mg, Fe, and Mn was done by standard methods using ICP–AES (Ca, Sr, Mg, Fe, Mn) and furnace AAS (Rb) on samples dissolved in 1.8 M acetic acid. The precision of the analysis was better than š10%. Ultrastructure was examined under gold coatings using a Zeiss DSM 940 SEM. To determine mineralogy, samples were analysed using a Phillips PW 1710 Diffractometer. Instrumental conditions were: CuKα radiation generated at 40 kV and 30 mA; scanning through 24– 32º 2Q at 0.5º=min, i.e. for the major diffraction peaks of aragonite, calcite, and dolomite. With these conditions the detection limit of our apparatus, determined by standard additions with pure phases, was about 0.2% calcite in aragonite and 0.5% dolomite in calcite. For 87 Sr=86 Sr analysis, cleaned macrofossil fragments were dissolved in 2.5 M hydrochloric or 1.8 M acetic acid. These preparations gave indistinguishable data. Pre-leached chalks were disaggregated with ultrasound in 2 ml of pure water and 30% of the sample was then dissolved by addition of 400 µl of 1.8 M acetic acid. After reaction, the solution was isolated for analysis by double centrifugation. Remaining undissolved sample was discarded. Values of 87 Sr=86 Sr were determined with a VG-354 fivecollector mass spectrometer using the multi-dynamic routines that include corrections for isobaric interference from 87 Rb [43]. Data have been normalised to a value of 0.1194 for 86 Sr=88 Sr. The data were collected between April, 1994, and September, 1997. During data collection, the measured value for NIST 987 was within 0.00002 of the value 0.710248. Data reported in Table 1 have been adjusted to a value of 0.710248 for NIST 987 which, in our laboratory corresponds to a value of 0.709175 for Modern Seawater Sr (MSS). Based upon replicated sample analysis the precision of our measurements (2 s.e.)
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was š15 ð 10 6 for single determinations, š11 for duplicates and š9 for triplicates. Total blanks were <2 ng of Sr. Sample contained >5 mg of Sr. Concentrations of Rb were found to be too low to require corrections for radiogenic Sr. Data from the literature that are reported here are adjusted to a value of 0.710248 for NIST 987.
6. Results and discussion 6.1. Sample integrity and diagenesis As all our samples have been buried at some time, they have been affected by diagenesis. We believe, for the reasons given below, that the alteration to 87 Sr=86 Sr is too small to affect our conclusions. Perhaps the most telling reason for our belief is that 87 Sr=86 Sr values in the K=T boundary (with values and 95% confidence intervals obtained by the non-parametric regression method given in [22]) are 0:707828 š 3 for Denmark and 0:707832 š 7 for Seymour Island, with the trend of the data from each locality being concordant (Figs. 1 and 2). Such good agreement, for different materials from opposite ends of the Earth, is unlikely to be fortuitous. It occurs only because each sequence is recording the original 87 Sr=86 Sr in K=T seawater. Other evidence supports this conclusion. Most importantly, macrofossils from Seymour Island retain their ultrastructural integrity, which suggests that they are effectively unaltered. Additionally, we have not analysed bulk chalk from Denmark; we have analysed about half of that part of a bulk sample that remains after an acid pre-leach has removed 30% of the sample (overgrowths, exchangeable Sr, acid-soluble clays). We expand on these points below and in the methods section. Firstly, chalk from northwestern Germany and eastern England yield excellent records of 87 Sr=86 Sr isotope evolution for the Late Cretaceous [39– 41,44], unquestionably record original 87 Sr=86 Sr values (when analysed by the method used here i.e. pre-leached) and are similar in appearance, chemical composition (Table 1), and diagenetic state (Fig. 3) to, and come from the same basin as, our Danish chalks. The effectiveness of our pre-leach method is clear from 87 Sr=86 Sr analysis of horizons of chalk at Hemmoor, northwestern Germany [44] in which
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Table 1 Chemical and isotopic data for K=T boundary sequences, Kjølby Gaard and Nye Klov, Denmark, and Seymour Island, Antarctica. Samples are macrofollis (SI) and acid-leached nannofossil chalks (KG, NK)
J.M. McArthur et al. / Earth and Planetary Science Letters 160 (1998) 179–192 Table 1 (continued).
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Fig. 3. SEM photomicrographs showing preservation of samples. (a) Chalk from Lagerdorf (sample 88 of [41]) that preserves a good record of original 87 Sr=86 Sr. (b) Chalk (KG12) from Kjølby Gaard. In (a) and (b) overgrowth is present in both but spar is rare. (c) and (d) Pycnodonte samples (c, 1175; d, 1163) from Seymour Island, showing unaltered compact layering in calcite. (e) Good preservation of Pseudotextularia deformis (long axis 200 µm) from 1 m below the K=T boundary, Kjølby Gaard. (f) Detail of aperture in (e) showing only minor nucleation of calcite on the specimen’s surface.
localised diagenetic cementation by carbonate has accompanied flint formation. Compared to nearby uncemented chalks, cemented horizons have 40%
lower Sr=Ca, but 87 Sr=86 Sr values (of pre-leached samples) that differ by less than 10 ð 10 6 (fig. 4 of [44]). Furthermore, samples of pre-leached nanno-
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fossil chalk (KG11) and macrofossil debris (KG12) from the same stratigraphic level in Kjølby Gaard have 87 Sr=86 Sr values that are analytically indistinguishable (Table 1). Finally, for twelve samples, we compared 87 Sr=86 Sr using pre-leaches that dissolved and discarded either 30% or 60% of the sample prior to analysis; the mean difference was 6 ð 10 6 , which suggests that our data are not significantly affected by diagenetic artefacts. The Seymour Island samples (all macrofossils) are monomineralic and have concentrations of Sr, Fe and Mn (Table 1) in the range shown by numerous authors ([3,45], and refs. therein) to be representative of well-preserved macrofossils. Ultrastructural detail is well preserved (Fig. 3). Replicate analyses on separate subsamples of macrofossil material show excellent agreement; this would not be expected were preservation poor. Finally, our macrofossils comprise a variety of faunal types, some retaining an original aragonitic mineralogy, whilst others retain their original calcite. Were these fossils altered, it would be unlikely that such different organisms would all alter to yield such similar isotopic ratios. Finally, samples considered poorly preserved by virtue of their mixed mineralogy, high Fe and Mn content, or anomalously high Sr content (Table 1) generally have an 87 Sr=86 Sr value much lower than that expected for the interval: in these sediments, alteration of macrofossil carbonate in our samples lowers 87 Sr=86 Sr rather than increasing it, as is usual. 6.2. Trends in 87 Sr=86 Sr When plotted against stratigraphic level (Figs. 1 and 2), 87 Sr=86 Sr changes by less than 20 ð 10 6 from the base (0.70784) to the top (0.70782) of the sequences. Statistical regression [22] of the data defines boundary values (with 95% CIs) as 0:707828 š 3 in Denmark and 0:707832 š 7 in Antarctica. Neither sequence shows positive boundary discontinuities in 87 Sr=86 Sr, as has been reported elsewhere for K=T sequences [4–8]. In both, however, there is a suggestion in the trends that values decline across the boundary and reach a minimum of 0.70782 just after it (about 0.5–1 m in Denmark; 20–40 m in Antarctica), thereafter increasing up the sequence through a broad maximum of 0.70783 before resuming the downward trend seen in the lower
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parts of the sequences. This trend is best seen in the Danish data, which are presumed to have a more extended post-boundary stratigraphic range than do the data from Antarctica. These data do not confirm the existence of a positive excursion in 87 Sr=86 Sr at the K=T boundary. A geological catastrophe, such as a meteorite impact, sudden climate change, sudden inundation of arid salt-laden continental areas, sudden glacial events, periods of acidic rain, etc., may well cause a sudden change in the 87 Sr=86 Sr of seawater, but, as correctly noted by Martin and Macdougall [6] and Nelson et al. [46], such excursions cannot suddenly disappear from seawater (and so the rock record) but must decay to the underlying trend in 87 Sr=86 Sr with a time constant that is the reciprocal of the residence time of marine Sr. Martin and Macdougall [6] point out that, with a residence time for Sr of 4 Ma, their putative spike should have decayed by only 20–25% of its value after 1 Ma, but that it had actually returned to pre-K=T values by that time, a puzzle they left unresolved (as they did the origin of the spike in 87 Sr=86 Sr of 45 ð 10 6 shown by their fig. 4 to occur about 2 Ma after the K=T boundary). It follows that any ‘spike’ with a duration of much less than 0.5 Ma must be an artefact of diagenesis or analysis or, if real, must decay exponentially to the background trend with a time constant that can reveal the true residence time of Sr. In Fig. 4a, we plot 87 Sr=86 Sr data against numerical age for our data and that of Martin and Macdougall [6], both superimposed on the modelled decay curve from a positive spike size of 28 ð 106 , the size reported by Martin and Macdougall [6]. The figure incorporates the underlying Maastrichtian increase of 19 ð 10 6 Ma 1 [47] and a Danian decrease of 15 ð 10 6 Ma 1 [22]. The residence time for Sr of 2 Ma was chosen to represent the shortest (and so fastest decay) thought reasonable, and that of 0.5 Ma to show the insensitivity of our conclusions to extreme choices. We have assigned numerical ages to the Seymour Island samples using an 87 Sr=86 Sr boundary value of 0.707832, a K=T boundary date of 65.0 Ma [51], and the fact that 1400 m of Maastrichtian sediments were deposited on Seymour Island and Snow Hill Island (A. Crame, pers. commun., 1997) during its 6.3 Ma [52] duration. We assume that sedimentation rate was constant
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Fig. 4. (a) Comparison of data with modelled 87 Sr=86 Sr decay curves. The model assumes an instantaneous spike of C28 ð 106 in 87 Sr=86 Sr at the K=T boundary [6], residence time for Sr of either 2 or 0.5 Ma and exponential decay of the 87 Sr=86 Sr spike according to the formula [87 Sr=86 Sr] D [87 Sr=86 Sr]o exp. t=RT/, where t D elapsed time, RT D oceanic residence time of Sr, and [87 Sr=86 Sr]o is the spike magnitude. Decay is superimposed on background trends for the late Maastrichtian of C0.000019 Ma 1 [47] and 0.000015 Ma 1 for the Danian [22]. Filled triangles are Seymour Island; filled circles are Denmark; open symbols from Martin and Macdougall [6], numerical ages for Sites 528 and 577 have been adjusted to include boundary hiatuses of 60 ka not recognised by them but reported elsewhere [48]. Data normalised to 0.710248 for SRM (NIST) 987. (b) Compilation of 87 Sr=86 Sr data as a function of numerical age for K=T samples reported here (filled circles) and those from Hess et al. [4] (hourglasses), Martin and Macdougall [6] (open circles), McArthur et al. [41] (open squares), McLaughlin et al. [44] (crosses), Sugarman et al. [47] (open triangles), Dennison et al. [49] (crossed squares), DePaolo and Ingram [50] (inverted triangles). All numeric ages have been adjusted to a common timescale where the K=T boundary is 65.0 Ma. Data of Hess et al. [4] adjusted using a value for their MSS of 0.709185, computed from their table 2 after exclusion of an obvious flyer, rather than their incorrectly computed value of 0.709198. Data normalised to 0.710248 for SRM (NIST) 987, except for Hess et al. [4], which is normalised to 0.709175 for MSS.
through the 1400 m of the Maastrichtian sequence. To the Danian part of the Danish sequence, we use the age model for Nye Kløv given by Keller et al. [23]. To the Maastrichtian part of the sequence we use a value of 4 cm ka 1 , one midway between the mean rate of 9 cm ka 1 for the entire Maastrichtian at its thickest in the Danish Basin [Kaminski 49], but viewed as too high by Keller et al. [23], and the values for several abyssal sites during K=T times of 0.3 to 2.5 cm 1 ka 1 [6]. Fig. 4a shows that our data fail to conform to the modelled decay curve, but do fit the underlying Maastrichtian and Danian trends. The data of Martin and Macdougall [6] scatter more widely than do ours; this scatter compromises a loose correspondence of Martin and Macdougall’s Danian data [6] to the modelled decay curve. Inter-laboratory bias of some 17 ð 10 6 in excess of that which can be accounted for by normalisation to standards has been noted to occur between laboratories [3,53]; whether this affects our comparison is not known. Data for K=T boundary strata from elsewhere are sparse, but are generally concordant with ours (Fig. 4b). Data for a Maastrichtian interval at Bidart, southwestern France [46] showing a boundary value of 0.707842 at the K=T boundary is in question after the discovery that foraminifera from this sequence have shells that luminesce under CL and contain high concentrations of iron (H. Vonhoff, pers. commun., 1996), facts that suggest to us that they have been altered despite preserving good external morphologies. A boundary sequence at Braggs, Alabama [54] shows a broad maximum, but no spike, through about 4 m of section centred on the boundary (boundary value about 0.70784 derived from fig. 1 of [54], corrected for interlaboratory bias [55]. The Sr-isotope regression of Sugarman et al. [47] assumes 87 Sr=86 Sr increases linearly to the boundary; their regression gives a value at 65.1 Ma of 0.707845 (at our postulated Maastrichtian maximum) but shows no discontinuities in the few data that define their trend prior to it (Fig. 4b for data points). Extrapolated to 65.0 Ma using a post-maximum decrease in 87 Sr=86 Sr of 15 ð 10 6 , their data predict a value of 0.707844 for the K=T boundary, a similar result to that given by these authors for a boundary sample from El Kef (0.707847 at 65.1 Ma).
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6.3. An explanation for spikes We believe that spikes and step functions in existing 87 Sr=86 Sr compilations for K=T boundary intervals result from analytical and diagenetic effects. The original report of a boundary spike, of 100 ð 10 6 , was made by Hess et al. [4] on the basis of two anomalous values (Fig. 4b). Neither measurement was duplicated; without such duplication neither data can be relied upon. The samples were 29-3-5-7 at Site 356, which these authors omitted to categorise for preservational quality (their table 2), and sample 12-5-19-21 from Site 577, for which there was no evidence of recrystallisation. Martin and Macdougall [6] report a maximum as occurring about 60 ka after the boundary, but they clearly state that preservation of samples was different on either side of the boundary at all the localities they examined. For example, for Site 356, their most densely sampled locality and the only one with a convincing discontinuity across the boundary, they state that “Maastrichtian samples frequently contain calcite overgrowths on their interior walls, while Danian foraminifera exhibit excellent preservation.” Diagenesis is expected to have affected boundary sediments at their sites as all have a hiatus at the boundary of at least 50 ka 1 [6,48]. The discontinuity in 87 Sr=86 Sr at Site 356 may result from diagenetic lowering of 87 Sr=86 Sr below the boundary. Meisel et al. [8] report a spike of 130 ð 10 6 across the K=T boundary at Sumbar, Turkmenistan. To isolate Sr from their samples, on the boundary clay itself, they were attacked with 5 M acetic acid without being pre-leached to remove contaminant Sr. Contrary to the statement of Meisel et al. that “leaching with 5 M acetic acid leaves clay minerals unaltered”, it leaves them leached of some of their radiogenic Sr; this treatment minimises, but does not prevent, dissolution of contaminant Sr from clay minerals by the hydrogen ion [3]. It is not surprising that their 87 Sr=86 Sr values appear to be lithology dependent; for those lithologies that might reasonably be expected to give the least contaminated ratios, because their Sr comes mostly from carbonate (marlstone S9, 13 to 16 cm above the boundary; marlstone S43, 0 to 1 cm below it), the mean ratio is 0:707834 š 10, a value close to our boundary values.
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The spike in 87 Sr=86 Sr reported by Meisel et al. is an analytical artefact. Attempts to use acids to isolate Sr from carbonate in clay-rich sediments are unlikely to succeed. Finally, Vonhoff and Smit [7] also report a boundary spike in 87 Sr=86 Sr around the K=T boundary. Most of their samples are rich in iron and manganese and luminesce brightly under CL, facts which compromise the reliability of their data by suggesting that their samples are altered. Those that are not apparently altered (two visually well preserved foraminifera from the E-clay, Guelhemmerburg Quarry) contain enough diagenetic carbonate (up to 1%) to yield the spike they report, even on the most conservative estimates of the Sr abundance and isotopic ratio in that diagenetic calcite, given that the foraminifera have resided for 65 Ma in a clay-rich (i.e. radiogenic) unit with a bulk 87 Sr=86 Sr of 0.75 (H. Vonhoff, pers. commun., 1998). The 87 Sr=86 Sr of Sr in seawater stopped increasing and started decreasing close to the K=T boundary [4,6,2,3,22]. Such a reversal requires a major change in the balance of Sr supplied to the ocean by mid-ocean ridges and rivers. We agree with the speculation of Vonhoff and Smit [7] that this decrease may have coincided with Deccan Trap volcanism and its consequences, but note that the downturn in 87 Sr=86 Sr at this time marked the end of a period of increase that lasted 25 Ma and the beginning of a period of decrease that lasted for 13 Ma [22]. The underlying reason for the change must therefore relate to mantle adjustments to the outpouring of the Deccan basalts, rather than the consequences of their surface weathering. Unfortunately, the number of variables, and the lack of quantification of their magnitudes, needed to construct models that address why such changes occur in the marine 87 Sr=86 Sr record preclude a useful discussion of this issue, a conclusion broadly in agreement with that of a few others [56,57].
7. Conclusions The mean 87 Sr=86 Sr at the K=T boundary, as determined from our combined data using non-parametric regression [22], is estimated as 0:707830 š 4 (95% CI), which is close to values found by others
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for this boundary. Our data do not show a spike or a step function in 87 Sr=86 Sr at the K=T boundary. The K=T boundary discontinuities in 87 Sr=86 Sr that have been reported elsewhere are artefacts of diagenesis or analytical methodology, or both. Marine 87 Sr=86 Sr peaked at or before 65.1 Ma and decreased thereafter into the Danian.
Acknowledgements The Radiogenic Isotope Laboratory at RHBNC is supported, in part, by the University of London as an ´ .M. McLaughlin intercollegiate facility. We thank O for assisting with the isotopic measurements, A. Osborn for assistance with sample preparation and elemental analysis, T. Styles and C. Stuart for help with preparation of figures and J. Davy for the SEM work. Samples from Kjølby Gaard were provided by A.S. Gale. We thank Finn Surlyk, Gerta Keller, Donald DePaolo, and two anonymous reviewers for constructive critiques of the script. [FA]
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