The influence of extraterrestrial material on the late Eocene marine Os isotope record

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ScienceDirect Geochimica et Cosmochimica Acta 144 (2014) 238–257 www.elsevier.com/locate/gca

The influence of extraterrestrial material on the late Eocene marine Os isotope record Franc¸ois S. Paquay a,⇑, Greg Ravizza a, Rodolfo Coccioni b b

a Dept. of Geology and Geophysics, University of Hawaii at Manoa, 1680 East West Road, POST 712, Honolulu, HI 96822, USA Dipartimento di Scienze della Terra, della Vita e dell’Ambiente dell’Universita, Campus Scientifico, Localita` Crocicchia, 61209 Urbino, Italy

Received 11 April 2013; accepted in revised form 21 August 2014; available online 1 September 2014

Abstract A reconstruction of seawater 187Os/188Os ratios during the late Eocene (36–34 Ma), based upon bulk sediment analyses from the sub-Antarctic Southern Atlantic Ocean (Ocean Drilling Program (ODP) Site 1090), Eastern Equatorial Pacific Ocean (ODP Sites 1218 and 1219) and the uplifted (land-based) Tethyan section (Massignano, Italy), confirms that the previously reported abrupt shift to lower 187Os/188Os is a unique global feature of the marine Os isotope record that occurs in magnetochron C16n.1n. This feature is interpreted to represent the change in seawater 187Os/188Os caused by the Popigai impact event. Higher in the Massignano section, two other iridium anomalies previously proposed to represent additional impact events do not show a comparable excursion to low 187Os/188Os, suggesting that these horizons do not record multiple large impacts. Comparison of records from three different ocean basins indicates that seawater 187Os/188Os begins to decline in advance of the Popigai impact event. At Massignano this decline coincides with a previously reported episode of elevated 3He flux, suggesting that increased influx of interplanetary dust particles (IDPs) contributed to the pre-impact shift in 187Os/188Os and not to the longer-term latest Eocene 187Os/188Os decline that occurred 1 million year after the Popigai impact event. Published by Elsevier Ltd.

1. INTRODUCTION Toward the end of the Eocene Epoch, multiple impact events occurred (Koeberl, 2009) contemporaneously with an episode of increased accretion of IDP’s (Farley, 2009). The cause of this episode of increased extraterrestrial (ET) influx is controversial. It is argued to result from either collisions in the asteroid belt of a L-type ordinary chondrite based on platinum group element (PGE) concentration in Popigai impact melts (Tagle and Claeys, 2005), or from a comet shower based on 3He flux in the Massignano section (Farley et al., 1998). More recently, Kyte et al. (2011) use chromium isotopes in late Eocene ejecta to show that the

⇑ Corresponding author.

E-mail address: [email protected] (F.S. Paquay). http://dx.doi.org/10.1016/j.gca.2014.08.024 0016-7037/Published by Elsevier Ltd.

ejecta contained extraterrestrial Cr indistinguishable from ordinary chondrites. Although this result is consistent with the conclusion of Tagle and Claeys (2005), Kyte et al. (2011) note that the Popigai PGE data (Tagle and Claeys, 2005) are also consistent with a H-type ordinary chondrite impactor. Therefore, Kyte et al. (2011) have speculated that the Branga¨ne asteroid family could be the source of the Popigai projectile. The suggestion is consistent with available data because the Branga¨ne asteroid family displays an infrared spectrum consistent with ordinary chondrites. These workers note that this includes the H-type chondrites, a meteorite group with cosmic-ray exposure age that clusters around 33–36 Ma, similar to the Popigai impact event. In a previous study Paquay et al. (2008a) used the impact-induced excursion in the late Eocene portion of the marine 187Os/188Os ratio record to calculate the amount of meteoritic Os that dissolved in seawater following the Popigai impact, and estimate the size of the projectile. Here,

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we report a new high-resolution late Eocene Os isotope record from the land-based Massignano section (Odin and Montanari, 1988; Premoli Silva and Jenkins, 1993) and expanded Os isotope records from three previously studied deep-ocean drilling sites. These records, together with complementary Os and Ir concentration data, are used to, (1) critically evaluate claims that two separate Ir anomalies in the Massignano section represent additional late Eocene impact horizons (Montanari et al., 1993; Bodiselitsch et al., 2004), (2) provide an updated discussion of impact-induced Os isotope excursions and projectile size estimation and (3) examine the relationship between the late Eocene episode of increased IDP flux (Farley et al., 1998) and the marine Os isotope record. These results have important implications for understanding the ways in which ET matter delivered to the Earth system influence the marine Os isotope record. 2. SAMPLE MATERIAL The sites studied in this contribution include the Massignano section (Coccioni et al., 1988: 43°320 1300 N; 13°350 3600 E), ODP Site 1218 from the Equatorial Pacific (8°53.3780 N, 135°22.000 W; 4826 m water depth: Lyle et al., 2002), the Southern Atlantic Ocean ODP Site 1090 (42°54.80 S, 8°54.00 E; 3700 m water depth: Gersonde et al., 1999) and a second Equatorial Pacific Site ODP 1219 (7°48.019N, 142°00.9400 W; 5063 m water depth) (Fig. 1). Massignano, ODP Site 1218 and Site 1090 were selected for further investigation because all three sites can be correlated to one another based on magnetostratigraphy, and previous works (Ravizza and Peucker-Ehrenbrink, 2003; Dalai et al., 2006; Paquay et al., 2008a) have established

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that these sediments effectively record secular change in the 187Os/188Os of seawater. The additional data reported here allow detailed comparison of late Eocene Os records from different oceanic basins. 2.1. The Massignano section Previously reported late Eocene Os isotope data from the Massignano section compare favorably with the Os records from multiple deep sea sediment records of similar age (Ravizza and Peucker-Ehrenbrink, 2003), but this original low-resolution dataset does not allow identification of any late Eocene impact. Here, we report a high resolution Os isotope record (110 new analyses) in the Scaglia Variegata Formation from 0.5 msl to about 6.4 msl, (meters stratigraphic level, where greater height corresponds to younger sediments). Additional samples between 10 and 11 msl were analyzed to further investigate earlier reports of Ir enrichment at this level (Bodiselitsch et al., 2004). These new data combined with previously published data (Ravizza and Peucker-Ehrenbrink, 2003) fall within the magnetochrons C16n.2n-C13r (Jovane et al., 2007), planktonic foraminiferal Zones P15–P16 or E14–E15 (Coccioni et al., 1988, 2009) and calcareous nannoplankton Zones NP18–NP19/ 20 (Coccioni et al., 1988; Monechi et al., 2000). The sediments that comprise the Massignano section were deposited at an estimated paleodepth of 1000–1500 m (Coccioni and Galeotti, 2003). They consist of marly limestone, with CaCO3 content varying from 50 to 80 wt% (average 76 wt%, Jovane et al., 2009). Sediment color alternates from greenish to reddish, likely due to changes in redox conditions on the sea floor. Biotite-rich layers of volcanic origin at 2.04, 5.25, 5.8, 6.5, 7.25, 7.75, 12.1 and 12.7 msl (Jovane

Fig. 1. Approximate locations of the studied ODP cores/uplifted section (squares) and impact craters (circles) of late Eocene age in a paleogeographical reconstruction of 35 Ma. The Popigai crater is the likely source of the global ejecta horizon within C16n.1n. Other known craters with ages that may overlap with the time interval studied here are indicated by numbers: Chesapeake Bay (1) Wanapitei (2) Mistastin (3) and Logoisk (4). The base image is taken from (http://jan.ucc.nau.edu/~rcb7/globaltext2.html).

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et al., 2009) were carefully avoided during sampling. Note that the magnetic stratigraphy at Massignano (Jovane et al., 2007) does not specify the depth of the boundary between magnetochrons C16n.2n and C16n.1r. To address this data gap we have placed this magnetic reversal boundary at 0 msl. The resulting age-depth relationship is very similar to the constant sedimentation rate model used in several earlier studies (Farley et al., 1998; Bodiselitsch et al., 2004). 2.2. ODP Site 1090 ODP Site 1090 (Hole B) consists of pale reddish-brown mud bearing diatom ooze and diatom-bearing nannofossil ooze. Over the studied interval the percentage of the carbonate and lithogenic sediment fraction average 20% and 66%, respectively (Anderson and Delaney, 2005) while organic carbon content is very low (0.05–0.1%) (Diekmann et al., 2005). Based on age constraints from previously reported paleomagnetic, biostratigraphic and chemostratigraphic data (Channell et al., 2003), our sampling resolution at this site is approximately one sample every 10 kyr. In this study, we augment published data (Paquay et al., 2008a), with 34 additional analyses from magnetochrons C15r and C15n, and 5 additional analyses from a 20 cm interval across the Popigai impact horizon within C16n.1n. 2.3. ODP Sites 1218–1219 ODP Sites 1218 and 1219 in the Eastern Equatorial Pacific are separated by 740 km and composed of clayey radiolarian ooze. During the late Eocene sediments at Site 1219 were deposited below the calcite compensation depth (CCD) while Site 1218 was located slightly above the CCD. Coherent Os isotope records from late Eocene sediments at both sites have been reported previously (Dalai et al., 2006; Paquay et al., 2008a) demonstrating that this difference in lithology does not affect the fidelity of the Os isotope record. At Site 1218, available data show that carbonate is variable (0.25–36 wt%) and total organic content is very low (0.02%) (Olivarez Lyle and Lyle, 2005). Site 1218 (Hole B) was investigated because the impact-induced Os isotope excursion was truncated in ODP 1219A due to a core-break and drilling disturbance immediately above the impact horizon (Paquay et al., 2008a). The record from ODP 1219A used here is composed almost entirely of previously published data (Dalai et al., 2006; Paquay et al., 2008a), with the exception of 3 additional analyses in the deep portion of the record. These samples were analyzed to ensure that the ODP 1219 record extended beyond 36 Ma. In this study data from ODP Sites 1218 and 1219 are combined to create a single composite Os isotope record that extends from below the Popigai impact horizon across the Eocene–Oligocene boundary. Fifty-two new Os analyses from ODP 1218B are combined with previously reported data from ODP 1218A (Dalai et al., 2006) using the adjusted and corrected revised meters composite depth (rmcd) scale reported in Tables 14 and 16 of Westerhold

et al. (2012). For both ODP 1218 and ODP 1219 rmcd is a measure of the depth of a given sediment sample below seafloor, where the word composite indicates that multiple drill holes at a given drill site have been correlated and placed on a common depth scale. The composite Site 1218 record is correlated to Site 1219 rmcd by linearly interpolating between tie points given in Table 22 of Westerhold et al. (2012) with the addition of a single tie point selected to align maximum Ir concentration in ODP1219A and ODP 1218B (see Section 4). Establishing these correlations allows merging data from ODP 1218 with data ODP 1219A (Dalai et al., 2006; Paquay et al., 2008a) into a single record. In the Equatorial Pacific records examined here nearly all of the relevant magnetostratigraphic datums are present in ODP 1219, but they are not all recognized in ODP 1218 (Pa¨like et al., 2005). Consequently we rely on the composite depth model described above (Pa¨like et al., 2005; Westerhold et al., 2012) to infer the depth of reversal boundaries in ODP 1218 data. These two studies interpret the magnetic reversal stratigraphy of Site 1219 slightly differently. Pa¨like et al. (2005) place the top of magnetochron C16n.1n at 181.08 rmcd, while Westerhold et al. (2012) place this magnetic reversal boundary at 181.79 rmcd, reducing the vertical extent of C16n.1n by roughly 50%. 2.4. Correlation and age model Age-depth models for all three sites are based upon linear interpolation between magnetostratigraphic datums (Table 1). While this approach may yield small systematic errors in age within magnetochrons, it ensures robust correlation of the late Eocene impact between sites, where it occurs consistently in C16n.1n and is attributed to the Popigai impact event (Liu et al., 2009). Ages of the relevant magnetic reversal boundaries (from the top of C16n.2n to the base of C13n) have been recently proposed in an updated Geologic Time Scale (GTS 2012: Gradstein et al., 2012). A similar but independent orbitally tuned age model that is derived in part from ODP Sites 1218 has also been recently published (Westerhold et al., 2014). Finally, the orbitally tuned age model for the Massignano section (Brown et al., 2009) can be combined with the magnetostratigraphy of the Massignano section (Jovane et al., 2007) to calculate a third set of ages for the late Eocene to early Oligocene magnetic reversal boundaries. At the Eocene–Oligocene boundary all three calibrations are in excellent agreement, but there is an offset of as much as 500 kyr among these different age models for the top and bottom of C16n.1n, the chron that includes the late Eocene impact event. This offset is large compared to the duration of C16n.1n. Consequently, employing site specific orbitally tuned age models for either ODP 1218/1219 or Massignano is inappropriate because to do so would imply different ages for magnetic reversal boundaries at different sites, and result in age offsets for the late Eocene impact horizon between sites. For these reasons we rely upon the simpler and more robust approach of linearly interpolating between magnetic reversals.

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Table 1 Depths of paleomagnetic reversal boundaries and ages used to construct age models used in this study.

C12n(o) C13n(y) C13n(o) Eocene–Oligocene boundary C15n(y) C15n(o) C16n.1n(y) C16n.1n(o) C16n.2n(y) C16n.2n(o) C17n.1n(y)

Age (Ma)a

Massignano (msl)b

ODP 1090 (mcd)c

ODP 1219 (rmcd)d

31.03 33.16 33.71 33.91 35.00 35.29 35.71 35.89 36.05 36.70 36.97

N.A. N.A. 21.5 19 11.05 9.27 6.22 5.2 4.15 0 N.A.

N.A. N.A. N.A.

142.82 169.77 173.43

284 287 289 293.5 296 309 319.8

180.16 Core gap 181.08e 182.52 183.07 187.47 188.92

N.A. – not applicable to this study. a Ages from GTS 2012 (Gradstein et al., 2012). b Meters section level (msl). c Meters composite depth (mcd) places multiple drill holes at a given drill site on a common depth scales. d Revised meters composite depth (rmcd). e See text in site description and Fig. 4 for additional information on the depth of the magnetic reversal boundary.

The GTS 2012 (Gradstein et al., 2012) time scale gives ages of 35.71 Ma and 35.89 Ma for the top and bottom of C16n.1n, respectively. This age range is consistent with the best available date for the formation of the Popigai impact crater (35.7 ± 0.2 Ma: Bottomley et al., 1997). However, the use of GTS 2012 rather than other available calibrations of the magnetic polarity time scale is somewhat arbitrary because small offsets in absolute age do not significantly affect our interpretations. Therefore our choice of time scale should not be construed as an endorsement of this particular calibration. Regardless of time scale choice, calculated average sedimentation rates vary significantly between ODP Site 1090, Massignano and ODP Sites 1218/1219. At ODP Sites 1218 and 1219 they vary between 0.3 and 0.8 cm/kyr in magnetochrons C16n.1r to C15n. This is much slower than in ODP Site 1090 where sediment accumulation rates are higher in general, reaching up to 3.8 cm/kyr in Magnetochron C13r. Sediment accumulation rate in Massignano is slightly higher than in Sites 1218/ 1219, averaging 0.7 cm/kyr. 3. ANALYTICAL METHODS Samples of approximately 20 cc from the ODP sites and the Massignano section were dried in an oven at a temperature of 60 °C for at least 2 days. Dried samples were powdered using a glass mortar and pestle, which were cleaned with MQ water and dried in between samples. Samples were stored in plastic vials until further processing. For individual analyses, 5–10 g of sample powder were digested (1000 °C fusion) by NiS fire assay in order to pre-concentrate the PGEs. Prior to fusion, samples were weighed and spiked with a tracer solution enriched in 105Pd, 190Os, 191 Ir, and 198Pt to allow concentration determination by isotope dilution. Isotope ratio measurements were made using an Element2 ICPMS with Os isotope ratio measurements performed by sparging (Hassler et al., 2000), and isotope ratio measurements of Pt, Pd and Ir were made in a

separate ICPMS run in which residual sample solutions from sparging were introduced into the ICPMS by conventional solution nebulization (Ravizza and Pyle, 1997). Only Os and Ir data are reported here. High quality Pt data, also generated in this study, are not reported here in order to keep the manuscript more focused. In contrast, Pd isotope data frequently show evidence of spectral interferences and are not reported because they are not considered reliable. Re analyses were performed on two samples from ODP Site 1090 in order to assess the potential influence of in situ decay of 187Re since the time of sediment deposition on measured 187Os/188Os ratios. These concentrations were also determined by isotope dilution following ICPMS analysis closely following the methods used by Ravizza and Vonderhaar (2012). An in-house Os standard was analyzed every five to six samples during each run to monitor the reproducibility and assess Os carry-over between samples. The average 187 Os/188Os of these standards was 0.1079 ± 0.0029 (2 SD; n = 51). Twenty procedural fusion blanks were carried out during the course of this study. These blanks yielded an average of 0.64 ± 0.20 pg Os/g of reagents, and 1.5 ± 0.5 pg Ir/g. A total weight of 31.5 g of mixed fusion reagents was used in the NiS fire assay method for blank determinations. Argon gas blanks were measured every set of five to six Os isotope analyses to monitor potential carryover of osmium between analyses. 4. RESULTS New results from this study extend Os isotope sediment records before and after the Popigai impact horizon, spanning a time window from roughly 36 Ma to beyond 34.6 Ma at multiple sites. In addition, high resolution records from Massignano and ODP 1218 refine our knowledge of the Os isotope excursion associated with the Popigai impact, confirming the initial findings of Paquay et al. (2008a) based on records from ODP 1090 and ODP 1219.

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Table 2 Osmium isotope, Os and Ir elemental data for bulk sediment, Massignano section (Ancona, Italy). Samples MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS

0.50 1.00 1.10 1.20 1.30 1.40 1.50 1.50Rd 1.50R 1.70 1.80 1.90 2.00 2.40 2.70 2.75 3.00 3.20 3.35 3.45 3.50 3.55 3.60 3.60R 3.60R 3.80 3.80R 4.00 4.10 4.20 4.50 4.50R 4.98/5.02 5.00 5.02/5.06 5.05/5.10 5.10/5.14 5.10/5.14R 5.14/5.18 5.18/5.22 5.22/5.26 5.26/5.30 5.30 5.30/5.34 5.30/5.34R 5.34/5.38 5.34/5.38R 5.38/5.42 5.42/5.46 5.46/5.50 5.50/5.54 5.50/5.54R 5.50/5.54R 5.54/5.58 5.58/5.62 5.60/5.62 5.60/5.62 5.60/5.62R 5.60/5.62R 5.60/5.62R 5.62/5.64

Height msla 0.50 1.00 1.10 1.20 1.30 1.40 1.50 1.50 1.50 1.70 1.80 1.90 2.00 2.40 2.70 2.75 3.00 3.20 3.35 3.45 3.50 3.55 3.60 3.60 3.60 3.80 3.80 4.00 4.10 4.20 4.50 4.50 5.00 5.00 5.04 5.08 5.12 5.12 5.16 5.20 5.24 5.28 5.30 5.32 5.32 5.36 5.36 5.40 5.44 5.48 5.52 5.52 5.52 5.56 5.56 5.61 5.61 5.61 5.61 5.61 5.63

AGE (Ma)b

187

Os/188Os

2SEc

36.63 36.55 36.54 36.52 36.51 36.49 36.48 36.48 36.48 36.45 36.43 36.42 36.40 36.34 36.30 36.29 36.25 36.22 36.20 36.19 36.18 36.17 36.16 36.16 36.16 36.13 36.13 36.10 36.09 36.07 36.02 36.02 35.93 35.93 35.92 35.91 35.91 35.91 35.90 35.89 35.88 35.88 35.87 35.87 35.87 35.86 35.86 35.86 35.85 35.84 35.83 35.83 35.83 35.83 35.83 35.82 35.82 35.82 35.82 35.82 35.81

0.524 0.485 0.527 0.589 0.524 0.512

0.007 0.006 0.022 0.011 0.009 0.009

35 37 31 24 11 31

0.550 0.547 0.521 0.520 0.526 0.528 0.537 0.501 0.518 0.446 0.457 0.474 0.526 0.523 0.469 0.532 0.461 0.518 0.257 0.508 0.495 0.521 0.552 0.511 0.533 0.503 0.481 0.520 0.510 0.457 0.503 0.508 0.492 0.491 0.509 0.486 0.497 0.507 0.484 0.501 0.466 0.407 0.424 0.458 0.477 0.466 0.387 0.452

0.014 0.008 0.008 0.017 0.007 0.025 0.005 0.013 0.009 0.011 0.036 0.013 0.010 0.008 0.006 0.013 0.012 0.004 0.007 0.016 0.017 0.010 0.008 0.006 0.009 0.016 0.021 0.007 0.011 0.008 0.004 0.008 0.010 0.009 0.006 0.006 0.011 0.014 0.008 0.009 0.008 0.008 0.004 0.026 0.001 0.008 0.003 0.005

124 37 30 37 67 33 27 30 27 14 22 47 38 39 27 29 28 36 84 29 35 23 38 33 25 58 49 49 43 44 37 39 43 46 34 20 68 55 62 60 37 51 58 71 59 32 71 47

0.453

0.004

88

0.432

0.008

56

Os (pg/g)

Ir (pg/g) 44 39 46 33 33 99 40 39 56 81 69 56 239 60 27 120 98 48 92 73 87 127 49 75 75 44 46 72 67 108 85 77 95 49 55 50 77 52 57 137 169 54 74 57 59 62 70 98 46 67 112 88 87

86 69 102 (continued on next page)

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Table 2 (continued) Samples

Height msla

AGE (Ma)b

187

2SEc

Os (pg/g)

Ir (pg/g)

MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS MAS

5.63 5.65 5.65 5.65 5.65 5.65 5.66 5.68 5.70 5.72 5.72 5.74 5.76 5.76 5.76 5.76 5.78 5.79 5.81 5.83 5.83 5.85 5.86 5.88 5.90 5.92 5.94 5.96 5.98 5.98 6.00 6.02 6.02 6.04 6.04 6.05 6.07 6.07 6.09 6.11 6.13 6.15 6.17 6.17 6.17 6.20 6.20 6.24 6.28 6.28 6.32 6.32 6.36 6.40 10.00 10.10 10.20 10.25 10.25 10.30 10.40 10.50

35.81 35.81 35.81 35.81 35.81 35.81 35.81 35.80 35.80 35.80 35.80 35.79 35.79 35.79 35.79 35.79 35.79 35.78 35.78 35.78 35.78 35.77 35.77 35.77 35.76 35.76 35.76 35.75 35.75 35.75 35.75 35.74 35.74 35.74 35.74 35.74 35.73 35.73 35.73 35.73 35.72 35.72 35.72 35.72 35.72 35.71 35.71 35.70 35.70 35.70 35.69 35.69 35.69 35.68 35.17 35.16 35.14 35.13 35.13 35.12 35.11 35.09

0.443 0.321

0.003 0.016

108 98

0.357

0.018

96

0.311 0.343 0.324 0.309

0.005 0.008 0.010 0.005

138 96 92 92

0.332 0.309 0.328 0.395

0.005 0.004 0.025 0.012

63 89 77 69

129 415 400 401 462 415 351 424 244 198 199 134 142

0.373 0.392 0.389

0.012 0.003 0.040

69 110 93

0.397 0.423 0.414

0.016 0.018 0.011

64 53 46

0.413 0.390 0.421 0.434

0.010 0.007 0.011 0.006

42 54 56 57

0.431 0.432 0.428 0.324 0.447 0.419

0.004 0.003 0.010 0.017 0.010 0.006

67 60 55 67 60 60

0.409

0.009

75

0.440 0.431 0.454 0.442 0.473 0.462 0.414 0.460 0.445 0.459

0.005 0.003 0.004 0.010 0.023 0.008 0.003 0.006 0.029 0.013

72 86 65 58 55 57 63 60 48 47

0.461

0.007

66

0.460 0.449 0.428 0.380 0.351 0.459 0.442 0.419 0.426 0.400

0.009 0.009 0.006 0.013 0.012 0.012 0.007 0.006 0.016 0.011

50 38 47 46 59 48 63 64 59 60

5.62/5.64R 5.64/5.66 5.64/5.66R 5.64/5.66R 5.64/5.66R 5.64/5.66R 5.66/5.67 5.67/5.69 5.69/5.71 5.71/5.73 5.71/5.73R 5.73/5.75 5.75/5.77 5.75/5.77R 5.75/5.77R 5.75/5.77R 5.77/5.78 5.78/5.80 5.80/5.82 5.82/5.84 5.82/5.84R 5.84/5.85 5.85/5.87 5.87/5.89 5.89/5.91 5.91/5.93 5.93/5.95 5.95/5.97 5.97/5.99 5.97/5.99 5.99/6.01 6.01/6.03 6.01/6.03R 6.03/6.05 6.03/6.05R 6.05/6.06 6.06/6.08 6.06/6.08R 6.08/6.10 6.10/6.12 6.12/6.14 6.14/6.16 6.16/6.18 6.16/6.18R 6.16/6.18R 6.18/6.22 6.18/6.22R 6.22/6.26 6.26/6.30 6.26/6.30 6.30/6.34 6.30/6.34 6.34/6.38 6.38/6.42 10.00 10.10 10.20 10.25 10.25R 10.30 10.40 10.50

Os/188Os

137 178 162 125 82 108 97 58 187 86 51 70 89 58 58 85 56 59 76 59 69 64 97 200 65 71 72 59 128 57 25 62 98 53 67 64 67 43 62 48 86 69 55 56 70 62 (continued on next page)

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Table 2 (continued) Samples

Height msla

AGE (Ma)b

MAS MAS MAS MAS MAS MAS

10.70 10.70 10.80 10.80 10.90 11.00

35.06 35.06 35.04 35.04 35.02 35.01

a b c d

10.70 10.70R 10.80 10.80R 10.90 11.00

187

Os/188Os

2SEc

Os (pg/g)

0.432 0.266 0.442 0.408 0.349

0.005 0.008 0.008 0.013 0.003

82 113 57 50 101

Ir (pg/g) 75 56 145 52 40

Stratigraphic height is given in meters section level (msl). Ages are calculated by linear interpolation between datums given in Table 1. 2SE indicates 2 times the standard error of the mean measured ratio based on in-run counting statistics. “R” indicates replicate analysis.

4.1. The Massignano section The new data from the Massignano section (Table 2) display a clear excursion to low 187Os/188Os within C16n.1n, coincident with a significant increase in Ir concentration at 5.65 meters section level (msl) (Fig. 2), corresponding to the most prominent previously reported Ir anomaly in the Massignano section (Bodiselitsch et al., 2004 and references therein). Importantly, 187Os/188Os ratios are declining during C16n.1r, below the 5.65 msl ejecta horizon. While there are instances of low 187Os/188Os and elevated Ir elsewhere in this record, these are isolated analyses that are likely the result of powder inhomogeneity because replicate analyses yield results more similar to adjacent samples. For example, replicate analyses of sample powders from 3.80 msl and 10.80 msl (Table 2) provide clear examples of this phenomenon. In contrast, replicate analyses from the Popigai impact horizon (5.64–5.77 msl) yield consistently low 187Os/188Os and high Ir concentrations. Even excluding outlying data at 3.80 and 10.80 msl with very low 187Os/188Os from the Massignano data set, the average relative standard deviation of replicate analyses is large (5% for n = 19) compared to other similar data sets (See ODP 1090 below for example). While Os concentrations also display a maximum at the 5.65 msl impact horizon within C16n.1n, absolute Os concentrations are roughly 3-fold lower than Ir concentrations, and are not easily differentiated from background Os concentrations. Note that plotted and tabulated Os isotope data have not been corrected for in situ decay of 187Re to 187Os because previous work in the Massignano section have yielded low Re concentrations (17–58 pg/g: Ravizza and PeuckerEhrenbrink, 2003) implying decay corrections to measured 187 Os/188Os ratios of less than 2%.

C15r and C15n. Five additional samples within the impact horizon were also analyzed to ensure that the magnitude and shape of the Os isotope excursion was not aliased by under-sampling. None of the plotted Os isotope data have been corrected for in situ Re decay. To evaluate the likely magnitude of in situ decay of 187Re to 187Os, we measured the Re content of two samples from site 1090 (samples 1090B-30X-1W-105/107; 1090B-31X-1W-107/109). Resulting Re concentrations are 0.18 and 0.10 pg/g, and corresponding 187Re/188Os ratios are 28 and 11, respectively. These results suggest that age corrections to measured 187 Os/188Os ratios at ODP 1090 are 0.02 or less, which is small compared to the dynamic range of the data (0.6 to 0.24). Down-core variation in measured 187Os/188Os at ODP 1090 display less scatter than in the Massignano section; the average relative standard deviation of replicate analyses (2% for n = 11including both new data and previously published data (Paquay et al., 2008a) is less than half of that documented in the Massignano data. The ODP 1090 Os isotope record shows that 187Os/188Os ratios begin to decline at 294 mcd, roughly 3 m below the primary impact layer, and remain below pre-impact values over the entire length of the studied interval (Fig. 3). These variations are directly correlated with Ir concentration variations, which display a gradual increase and then a single welldefined maximum. The Os isotope record shows great coherence in spite of large variations in Os concentration. As in the Massignano record, Os concentration displays a local maximum at the impact horizon that is small compared to the Ir maximum. In spite of a local maximum in Os concentration at the impact horizon, the Os concentration profile does not provide a clear indication of the impact horizon when considered separately from other data, such as Ir.

4.2. ODP Site 1090 4.3. ODP Sites 1218–1219 Elevated Ir concentrations (Kyte, 2001) and an excursion to lower 187Os/188Os (Paquay et al., 2008a) are known to coincide with microtektites and clinopyroxene (cpx) spherules from the Popigai impact within magnetochron C16n.1n at 291 mcd at ODP Site 1090 (Liu et al., 2009) (Fig. 3). The previously reported late Eocene Ir–Os isotope data from ODP Site 1090 (Paquay et al., 2008a) has been augmented with new data (Table 3) that extends the record an additional 4 m up-section to include magnetochrons

Analyses of ODP 1218B were performed to record the post-impact 187Os/188Os recovery in the Equatorial Pacific and a few samples from ODP 1219A were performed extending this record back to roughly 36 Ma at both Sites 1218 and 1219 (Table 4). These new results are combined with previously published data (Dalai et al., 2006; Paquay et al., 2008a) to provide a composite record spanning magnetochrons C16r.2n through C13r (Fig. 4). Plotted

F.S. Paquay et al. / Geochimica et Cosmochimica Acta 144 (2014) 238–257

Fig. 2. Bulk sediment 187Os/188Os ratios (upper panel), Os concentration (middle panel) and Ir concentrations (bottom panel) in the Massignano section shown in stratigraphic height. All data are from this study. Magnetochrons are from Jovane et al. (2007). Ages from Gradstein et al. (2012) are shown. Individual replicate analyses are shown as open circles. Filled circles show the full data set, both individual analyses and the average of all replicate analyses at a given depth. Error bars indicate the standard deviation associated with each set of replicate analyses. All data are from this study. 187

Os/188Os data have not been corrected for the decay of Re to 187Os because slowly accumulating pelagic sediments tend to have extremely low Re/Os ratios. In addition, previous work (Dalai et al., 2006) has shown good agreement between measured 187Os/188Os from these two sites 187

245

Fig. 3. Bulk sediment 187Os/188Os ratios (upper panel), Os concentration (middle panel) and Ir concentrations (bottom panel) in ODP Site 1090. Open circles are data from this study. Open squares are from Paquay et al. (2008a). Magnetochrons are from Channell et al. (2003). Ages are from Gradstein et al. (2012).

across the Eocene–Oligocene transition, indicating that the influence of 187Re decay on 187Os/188Os ratios is small enough to ignore. One initial objective of this study was to use data from ODP 1218B to fill gaps in the record from ODP 1219A. (See Funakawa et al., 2006 for additional discussion of gaps in the ODP 1219 record.) Recent reinterpretation of the magnetic stratigraphy in ODP 1219A (Westerhold et al., 2012) is consistent with truncation of magnetochron C16n.1n in ODP 1219A by a hiatus, because the top of

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Table 3 Os and Ir concentration and

187

Os/188Os data from ODP Site 1090 (Southern Atlantic Ocean).

Samples

Depth (mcd)a

Age (Ma)b

187

2SEc

Os (pg/g)

1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B 1090B

284.14 284.24 284.33 284.44 284.54 284.64 284.82 284.94 285.04 285.14 285.34 285.44 285.63 285.73 285.83 285.94 286.05 286.23 286.33 286.45 286.54 286.64 286.64 286.84 287.05 287.23 287.42 287.63 287.87 287.92 288.04 288.13 288.23 288.43 290.95 290.99 291.04 291.06 291.12

35.01 35.02 35.03 35.04 35.05 35.06 35.08 35.09 35.10 35.11 35.13 35.14 35.16 35.17 35.18 35.19 35.20 35.22 35.23 35.24 35.25 35.26 35.26 35.28 35.30 35.34 35.38 35.42 35.47 35.48 35.51 35.53 35.55 35.59 35.79 35.79 35.79 35.79 35.79

0.490 0.490 0.510 0.501 0.594 0.481 0.457 0.465 0.470 0.452 0.450 0.564 0.443 0.445 0.436 0.443 0.445 0.444 0.413 0.441 0.380 0.438 0.434 0.443 0.429 0.442 0.449 0.434 0.428 0.466 0.427 0.415 0.413 0.407 0.283 0.273 0.270 0.250 0.273

0.003 0.005 0.008 0.005 0.004 0.003 0.002 0.002 0.011 0.003 0.004 0.020 0.006 0.008 0.007 0.002 0.002 0.004 0.004 0.005 0.004 0.003 0.003 0.003 0.006 0.005 0.007 0.001 0.003 0.007 0.003 0.004 0.003 0.003 0.002 0.001 0.002 0.001 0.001

64 51 49 77 93 64 37 54 33 27 42 64 42 42 49 54 60 52 66 66 66 91 72 108 41 84 64 60 88 77 93 71 69 72 150 185 177 240 152

a b c d

30X 1W 15/17 30X 1W 25/27 30X 1W 34/36 30X 1W 45/47 30X 1W 55/57 30X 1W 65/67 30X 1W 83/85 30X 1W 95/97 30X 1W 105/107 30X 1W 115/117 30X 1W 136/137 30X 1W 145/147 30X 2W 14/16 30X 2W 24/26 30X 2W 34/36 30X 2W 44/46 30X 2W 55/57 30X 2W 74/76 30X 2W 84/86 30X 2W 94/96 30X 2W 104/106 30X 2W 114/116 30X 2W 114/116Rd 30X 2W 134/136 30X 3W 6/7.5 30X 3W 24/26 30X 3W 43/45 30 3W 64/66 30X 3W 88/90 30X 3W 93/95 30X 3W 105/107 30 3W 114/116 30X 3W 124/126 30X 3W 144/146 30 5W 97/98 30X 5W 103/104 30X 5W 105/106 30X 5W 107/109 30X 5W 113.5/114.5

Os/188Os

Ir (pg/g) 46

60 30 35 17 12 7 94

101 34 30 54 84 37 95 138 62 130 37 196 114 99 59 86 715 767 814 1118

Depths for ODP1090 samples are given in meters composite depth (mcd). Ages are calculated by linear interpolation between datums given in Table 1. 2SE indicates 2 times the standard error of the mean measured ratio based on in-run counting statistics. “R” indicates replicate analysis.

C16n.1n nearly coincides with the Ir maximum and 187 Os/188Os minimum in this core (Fig. 4). This differs from Massignano (Fig. 2) and ODP 1090 (Fig. 3) where the impact horizon is roughly centered in magnetochron C16n.1n. In addition, the Ir maximum in ODP 1218B is not at the rmcd depth predicted by either Pa¨like et al. (2005) or Westerhold et al. (2012) based on correlation of physical properties between these two sites. Recognition of a small gap in ODP 1218B does not affect our earlier conclusion (Paquay et al., 2008a) that the Ir and Os isotope excursions in ODP 1219A occur within magnetochron C16n.1n and correspond to the Popigai impact event. However, we cannot be certain that the composite ODP 1218/ 1219 record provides a continuous record of Os isotope variations from C16n.2n through C15n.

In spite of possible gaps in the composite ODP 1218/ 1219 record, several features are robust. Prior to the late Eocene impact within magnetochron C16n.2n, 187Os/188Os ratios are slightly higher than 0.5 and display relatively little variation, particularly in ODP 1219A where Os concentrations are large and variable. The very high Ir concentrations (maximum values of 3.7 ng/g in ODP 1218B and 1.6 ng/ g in ODP 1219A) and corresponding low 187Os/188Os (minimum values of 0.24 in ODP 1218B and 0.31 in ODP 1219A) give us confidence in correlating data from these two sites in order to align Ir maxima (Fig. 4). The length scale of the recovery from the impact event in ODP 1218B is quite short, as indicated by the very rapid decrease in Ir concentration immediately above the Ir maximum and the shift from minimum 187Os/188Os of 0.24 to a plateau at

F.S. Paquay et al. / Geochimica et Cosmochimica Acta 144 (2014) 238–257 Table 4 Os and Ir concentration and

187

Os/188Os data from ODP Sites 1218 and 1219 in the Eastern Equatorial Pacific.

Sample

1218 rmcda

1219 rmcdb

Age (Ma)c

1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B 1218B

239.90 240.66 240.80 245.14 245.61 245.67 245.67 245.73 245.81 245.87 245.93 246.01 246.07 246.13 246.13 246.21 246.21 246.27 246.32 246.37 246.43 246.51 246.51 246.57 246.63 246.63 246.71 246.77 246.83 246.89 246.97 246.97 247.05 247.11 247.17 247.23 247.31 247.31 247.37 247.43 247.46 247.50 247.56 247.56 247.62 247.70 247.76 247.80 247.86 247.92 248.05 248.12 248.20 248.26 248.32 248.38 248.46 248.60 248.66

173.76 174.29 174.38 178.77 179.24 179.30 179.30 179.35 179.43 179.49 179.55 179.63 179.69 179.75 179.75 179.83 179.83 179.88 179.93 179.98 180.04 180.11 180.11 180.16 180.20 180.20 180.27 180.32 180.37 180.41 180.48 180.48 180.54 180.59 180.64 180.69 180.75 180.75 180.80 180.85 180.88 180.91 180.96 180.96 181.01 181.07 181.12 181.15 181.20 181.25 181.35 181.41 181.47 181.52 181.57 181.62 181.68 181.79 181.79

33.77 33.87 33.89 34.73 34.82 34.83 34.83 34.84 34.86 34.87 34.88 34.90 34.91 34.92 34.92 34.93 34.93 34.95 34.96 34.96 34.98 34.99 34.99 35.00 35.03 35.03 35.08 35.12 35.16 35.19 35.24 35.24 35.29 35.33 35.37 35.40 35.45 35.45 35.49 35.53 35.55 35.58 35.61 35.61 35.65 35.70 35.71 35.72 35.72 35.73 35.74 35.75 35.76 35.76 35.77 35.78 35.78 35.80 35.80

24X 1W 30/32 24X 1W 106/108 24X 1W 120/122 24X 4W 106/108 24X 5W 4/6 24X 5W 10/12 24X 5W 10/12Re 24X 5W 16/18 24X 5W 24/26 24X 5W 30/32 24X 5W 36/38 24X5W 44/46 24X 5W 50/52 24X 5W 56/58 24X 5W 56/58R 24X 5W 64/66 24X 5W 64/66R 24X 5W 70/72 24X 5W 75/77 24X 5W 80/82 24X 5W 86/88 24X 5W 94/96 24X 5W 94/96R 24X 5W 100/102 24X 5W 106/108 24X 5W 114/116 24X 5W 114/116R 24X 5W 120/122 24X 5W 126/128 24X 5W 132/134 24X 5W 140/142 24X 5W 140/142R 24X 5W 148/150 24X 6W 4/6 24X 6W 10/12 24X 6W 16/18 24X 6W 24/26 24X 6W 24/26R 24X 6W 30/32 24X 6W 36/38 24X 6W 40/42 24X 6W 44/46 24X 6W 50/52 24X 6W 50/52R 24X 6W 56/58 24X 6W 64/66 24X 6W 70/72 24X 6W 74/76 24X 6W 80/82 24X 6W 86/88 24X 6W 99/102 24X 6W 106/107 24X 6W 114/116 24X 6W 120/122 24X 6W 126/127 24X 6W 132/134 24X 6W 140/142 24X 7W 4/6 24X 7W 10/12

247

187

Os/188Os

2SEd

0.406 0.402

0.002 0.002

46 45

0.329 0.341

0.004 0.002

68 107

Os (pg/g)

Ir (pg/g) 28

0.379 0.368

0.006 0.003

66 70

0.394 0.454

0.004 0.006

58 54

0.397 0.417 0.450

0.008 0.009 0.009

55 75 65

0.519 0.405 0.457

0.007 0.001 0.003

54 58 79

0.454 0.460 0.500 0.491 0.483

0.002 0.008 0.004 0.005 0.004

59 37 32 64 52

0.481 0.483

0.003 0.012

34 105

0.477 0.502 0.534 0.505

0.003 0.004 0.015 0.006

33 84 45 45

0.503

0.002

68

0.487 0.490 0.501 0.497 0.542 0.408 0.490 0.449 0.487

0.007 0.010 0.009 0.005 0.013 0.002 0.003 0.002 0.005

80 66 62 150 36 74 65 36 34

0.470

0.004

39

107 387 158 316 216 222 149 778 233 146 370 302 142 158 113

143 132

192 37 192 142 235 206 214

165 188 117 333 241 139 150 161

0.485

0.006

0.451 0.450 0.475

0.003 0.005 0.003

47

24 387 258 287

35 41 285 103 279 (continued on next page)

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Table 4 (continued) Sample

1218 rmcda

1219 rmcdb

Age (Ma)c

187

Os/188Os

2SEd

Os (pg/g)

Ir (pg/g)

1218B 24X 7W 16/18 1218B 24X 7W 24/26 1218B 24X 7W 24/26R 1218B 24X 7W 30/32 1218B 24X 7W 36/38 1218B 24X 7W 36/38R 1218B 24X 7W 40/42 1218B 24X 7W 43/44 1218B 25X 1W 14/16 1218B 25X 1W 25/27 1218B 25X 1W 30/32 1218B 25X 1W 52.5/54.5 1218B 25X 1W 57/59 1218B 25X 1W 60/62 1218B 25X 1W 72/74 1218B 25X 1W 78/80 1218B 25X 1W 94/96 1218B 25X 1W 101/103 1218B 25X 1W 112/5/114.5 1218B 25X 1W 125/127 1219A 18H 3W 125/129 1219A 18H 3W 125/129R 1219A 018H 03W 140/142 1219A 018H 03W 140/142R

248.72 248.80 248.80 248.86 248.92 248.92 248.96 248.99 250.38 250.54 250.61 250.91 250.97 251.00 251.14 251.21 251.38 251.45 251.56 251.67

181.79 181.80 181.80 181.81 181.81 181.81 181.90 181.97 184.25 184.49 184.59 185.03 185.14 185.17 185.39 185.49 185.75 185.85 186.02 186.19 185.76 185.76 185.90 185.90

35.80 35.80 35.80 35.80 35.80 35.80 35.81 35.82 36.22 36.26 36.28 36.34 36.36 36.36 36.39 36.41 36.45 36.46 36.49 36.51 36.45 36.45 36.47 36.47

0.424 0.360

0.005 0.004

40 261

0.312 0.240 0.251 0.245 0.325

0.002 0.002 0.003 0.003 0.003

73 302 780 162 26

0.522 0.546 0.502

0.004 0.007 0.013

17 22 25

0.548

0.004

29

0.571 0.524

0.020 0.021

24 583

0.449 0.521

0.009 0.012

274 344

80 1454 1058 1789 2828 2735 3670 551 941 264 349 356 128 99 494 497 101 156 189 172 88 86 97 82

a

Depths for samples from ODP 1218B are given in adjusted rmcd (revised meters composite depth) from Westerhold et al. (2012). Depths for 1218B samples are correlated into the ODP 1219 rmcd scale as described in Westerhold et al. (2012). See section 2.3 for additional detail. c Ages are calculated by linear interpolation between datums given in Table 1. d 2SE indicates 2 times the standard error of the mean measured ratio based on in-run counting statistics. e “R” indicates replicate analysis. b

approximately 0.45 above the impact horizon over a length scale of 20 cm. The short length scale of the Os isotope recovery is consistent with the presence of a hiatus at the top of C16n.1n as proposed by Westerhold et al. (2012), and discussed above. As suggested by the previous two sites, Os concentration data do not serve as reliable indicator of the Popigai impact horizon. 5. DISCUSSION 5.1. Evaluating the number of impacts recorded in the Massignano section Investigation of the coupled variation of 187Os/188Os and the concentrations of Os and Ir in the Massignano section is motivated by previous suggestions that elevated Ir concentrations at approximately 6.2 and 10.3 msl correspond to two impact horizons (Montanari et al., 1993; Bodiselitsch et al., 2004) that are distinct from the well established impact horizon at approximately 5.61 msl (Montanari et al., 1993; Langenhorst, 1996; Clymer et al., 1996a; Pierrard et al., 1998). Bodiselitsch et al. (2004) noted that elevated Ir at the 6.17–6.19 msl interval is associated with elevated concentrations of Ni and Cr, siderophile elements that are also relatively enriched in most meteoritic material. These workers suggested that the two additional horizons with siderophile enrichments could be related to

known late Eocene impact craters (Chesapeake Bay, Mistastin, Wanapitei, and Logoisk impact craters), which have similar ages. More carefully evaluating the claim of multiple late Eocene impact horizons in the Massignano section is important as part of ongoing efforts to better understand the underlying cause of the cluster of late Eocene impacts (Koeberl, 2009), and their possible climatic effects (Bodiselitsch et al., 2004; Coccioni et al., 2009), and potentially for testing proposed calibrations of the late Eocene portion of the geologic time scale that are based on orbital tuning (Westerhold et al., 2014) as noted in Section 2.4. New Massignano Ir data spanning 0.5–6.4 msl and 10.0 11.0 msl reveal only a single interval (5.60–5.65 msl) of persistently elevated Ir concentrations (>200 pg/g: Fig. 2). This Ir peak reaches 462 pg/g (5.65 msl) compared to 199 pg/g at 5.61 msl measured by Montanari et al. (1993) and 280 pg/g at 5.62–5.68 msl (Pierrard et al., 1998). Differences in absolute Ir concentrations between studies may originate from analytical differences (neutron activation versus ICPMS in this study). Alternatively, heterogeneities resulting from variable proportions of Ir-rich cpx spherules in the sediment analyzed may also contribute to the observed concentration differences. This interval of persistently elevated Ir in the Massignano section overlaps with the presence of shocked quartz (Langenhorst, 1996; Clymer et al., 1996a), Ni-rich spinels and microcrystites (Pierrard et al., 1998). Coincident with the Ir anomaly,

F.S. Paquay et al. / Geochimica et Cosmochimica Acta 144 (2014) 238–257

Fig. 4. Bulk sediment 187Os/188Os ratios (upper panel), Os concentration (middle panel) and Ir concentrations (bottom panel) in ODP Site 1218A/B (filled symbols; this study) and 1219 (open symbols; Dalai et al., 2006; Paquay et al., 2008a, and this study). Ages are from Gradstein et al. (2012). Magnetostratigraphic datums are from Pa¨like et al. (2005). The dark gray band indicates the thickness of C16n.1n, as proposed by Westerhold et al. (2012). This is significantly reduced compared to that proposed by Pa¨like et al. (2005) which is shown by the width of the dark and light bands combined. See Sections 2.3 and 4.3 for additional discussion.

there is also a distinctive excursion to low 187Os/188Os over the 5.70–5.65 msl interval, reaching as low as 0.309 (Fig. 2). This pattern of elevated Ir and low 187Os/188Os within magnetochron C16n.1n was also observed at ODP sites 1219 and 1090 (Paquay et al., 2008a). At ODP Site 1090

249

the Os isotope excursion coincides with the late Eocene cpx-bearing spherule layer (Kyte, 2001). A careful review of characteristics and global distribution of this layer (Liu et al., 2009) makes a strong case that the source for this particular ejecta horizon is the Popigai impact crater, as previously suggested. For example, the similar Sr–Nd isotope signatures of the target rock and the cpx spherules provide strong evidence for this connection (Kettrup et al., 2003). The Massignano data from this study do not display consistently high Ir concentrations or pronounced excursions to low 187Os/188Os at depths other than the 5.70–5.65 msl interval within C16n.1n that corresponds to the Popigai impact event (Liu et al., 2009). In other proposed impact horizons at 6.17–6.20 msl and 10.20– 10.25 msl (Montanari et al., 1993; Bodiselitsch et al., 2004) there is no evidence of correlated excursions to high Ir and low 187Os/188Os. In the 6.17–6.20 msl interval one of our Ir concentration analysis (128 pg/g) is suggestive of Ir enrichment. This concentration is an intermediate value compared to previous studies (95 pg/g in Montanari et al., 1993); 259 ± 32 in Bodiselitsch et al. (2004), but two additional replicate Ir analyses (50–60 pg/g) of this same sample powder failed to reproduce the high Ir concentration. In addition, the measured 187Os/188Os (0.473 ± 0.023) of the sample with slightly elevated Ir is similar to the other replicate analyses (0.442 ± 0.010, 0.462 ± 0.008) and to surrounding samples. These Os isotope ratios are inconsistent with a significant meteoritic component. In the 10.0–11.0 msl interval, Ir concentrations from the present study range from 46 to 113 pg/g and do not show any particular increase in the 10.2–10.3 msl interval (48–64 pg/g; n = 4). This stands in contrast to previously reported Ir data (341 pg/g: Montanari et al., 1993; 149 pg/g: Bodiselitsch et al., 2004). In two samples adjacent to this interval, 187Os/188Os ratios are low, as low as 0.351, but a distinctive isotope excursion, comparable to that documented at 5.65 msl, is not present (Fig. 2). The 187Os/188Os and Ir records from ODP 1090 (Fig. 3) and ODP 1218/1219 (Fig. 4) do not give any indication of additional impact events that could be correlated with the proposed impact horizons at 6.2 and 10.3 msl in the Massignano section. If these horizons at Massignano were the product of significant impact events, then one might expect to find correlated Ir enrichments and excursions to low 187 Os/188Os at other sites around the globe. The 6.2 msl horizon at Massignano occurs in the latter part of C16n.1n which allows identification of time correlative intervals of ODP 1090 and ODP 1218/1219. Although this time interval may be absent from the ODP 1218/1219 record (see Section 4.3 above), it is present in ODP 1090. In the ODP 1090 record there are two Ir analyses late in C16n.1 that are clearly elevated compared to surrounding samples (Fig. 3). However, these are not adjacent samples in the core and they are not associated with clear excursions to low 187Os/188Os. Consequently, we conclude that there is no evidence of an impact correlative with the Massignano 6.2 msl horizon in ODP 1090. The situation for the Massignano interval at 10.3 msl, which occurs in C15n, is similar. Neither the relatively condensed (slow accumulation rate)

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Table 5 Comparison of soluble impact-derived Os and Ir fluence measurements at 4 different sites. Popigai impact event

187

187

Rpre-impact

Massignano ODP 1090C ODP 1218B ODP 1219A

0.47 0.47 0.47 0.47

Os/188Os

Os/188Os

Rpost-impact

f: Fraction Os(sw) from projectilea

M(Os)impactb (moles dissolved)

Ir fluence (ng/cm2)c

0.31–0.24d 0.250 0.240 0.305

0.47 0.65 0.68 0.48

0.63–1.5  108 1.3  108 1.5  108 0.66  108

4.5 15 60 >40

a The fraction of the total dissolved 188Os in seawater that is derived from the impacting projectile is calculated as f = (Rpost-impact  Rpre-impact)/(Rimpactor  Rpre-impact), where the pre-impact 187Os/188Os of seawater (Rpre-impact), the immediately post-impact and 187Os/188Os of seawater (Rpost-impact) are obtained from the marine Os record (e.g., Figs. 2–4). The 187Os/188Os of the all projectiles (Rimpactor) is assumed to be 0.13 for simplicity. b The quantity (moles) of projectile-derived Os dissolved in seawater immediately after an impact event is calculated as: M(Os)impact = [f/ [(1  f)] * (7.3  107 moles Os). Where the value of 7.3  107 corresponds to inventory of dissolved Os in the modern ocean in moles and the difference between f and the fraction of total impact-derived Os resulting from variations in the abundance of 188Os is ignored. c The sources of Ir fluence estimates are: Massignano and ODP 1090 are from Liu et al. (2009). In ODP 1218B the Ir fluence was estimated by linearly interpolating the fluence between data points between 20 and 42 cm depth in Section 7 of core 24X (n = 5) after subtracting a background value 300 pg/g and using a dry bulk density of 1 g/cm3. This values was then doubled because the bottom half of the Ir peak is truncated at the bottom of core 24X (Table 4). An analogous approach was used to estimate Ir fluence for ODP 1219 (Data from Paquay et al., 2008a). d Low value based on a single analysis at 5.6 msl (Clymer et al., 1996b).

ODP 1218/1219 record nor the relatively expanded ODP 1090 record show any sign of elevated Ir and associated low 187Os/188Os within C15n. Thus our data from additional sites show no evidence of impact horizons that correlate with either 6.2 or 10.3 msl at Massignano. Based on the results discussed in the preceding paragraphs we conclude that previous reports of elevated Ir concentrations at 6.2 and 10.2 msl are not best interpreted as the result of impact events. Rather, it is more likely that these instances of elevated Ir are the result of either authigenic enrichment of ambient dissolved Ir from seawater, or perhaps the local influence of discrete Ir-rich particles. Indications that Ir concentrations in excess of 200 pg/g need not reflect the presence of meteoritic material can be found elsewhere in the Massignano data set; see for example data from the 2.7 msl interval where Ir is 239 pg/g and 187 Os/188Os is 0.501 ± .013 (Table 2). In the Bodiselitsch study, (2004), Ir concentrations positively correlate with Co concentrations, an element whose inventory in pelagic marine sediment is almost entirely scavenged from seawater (Kyte et al., 1993). In pelagic sediments Ir and Co concentrations tend to correlate inversely with sedimentation rate (Barker and Anders, 1968; Kyte et al., 1993), thus the small enrichment in siderophile elements observed by Bodiselitsch et al. (2004) may simply reflect minor episodes of slow sediment accumulation. Alternatively, it is possible that previous studies (Montanari et al., 1993; Bodiselitsch et al., 2004) encountered discrete Ir-rich phases. As noted in the Massignano results section above, the presence of such phases can be inferred from disparities in replicate analyses at 10.8 and 3.8 msl. Although Schmitz et al. (2009) establish that the inventory of definitively extraterrestrial chromiumrich spinel grains (0.006 grains/kg sediment) does not indicate an episode of increase meteoritic flux during the late Eocene, this study also documents considerably larger quantities of other Cr-rich spinel grains (>2 grains/kg sediment). Grains such as these are also possible carries of high Ir concentrations. Although more than one process is likely

responsible for producing local Ir enrichments in the Massignano section, comparison of the Massignano record (Fig. 2) to the records from ODP 1090 (Fig. 3) and ODP 1218/1219 (Fig. 4) provides strong evidence that there is only one large impact event that is recorded globally within the portions of the Massignano section studied here. This event occurs within C16n.1n and is attributed to the Popigai impact event (Liu et al., 2009). 5.2. Popigai projectile size estimates and the impact-induced Os isotope excursion from multiple sites New Os isotope records of the Popigai impact event from Massignano and ODP 1218B, as well as additional analyses of the impact horizon at ODP 1090B, are fully consistent with our initial work proposing the use of the marine Os isotope record to estimate projectile size (Paquay et al., 2008a). Os-based estimates of projectile size are important for two main reasons. First, in the absence of an impact crater, the marine Os isotope record is very likely the best available means of estimating the size of siderophile-rich projectiles. Second, for cases in which craters are identified, the computational models of impact crater formation that are used to estimate projectile size, velocity, and angle of incidence still rely heavily upon geochemical data such as Ir fluence, to select the most likely set of projectile attributes (Artemieva and Morgan, 2009). Impactinduced excursions in the marine Os isotope record have potential to be used in a similar fashion. The Os isotope excursion associated with the Popigai impact yields four independent estimates of the amount of impact-derived Os that dissolved in seawater (M(Os)impactor, expressed in moles). This is the key parameter in Os-based estimates of projectile size (Table 5). These estimates of M(Os)impactor range from 0.6 to 1.5  108 moles Os and are consistent with the earlier estimate of 1.1  108 moles Os (Paquay et al., 2008a). The approach for estimating M(Os)impactor assumes two-component

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251

Table 6 Summary of all know impact-induced excursions in the marine Os isotope record and associated craters. Event

Crater diameter (km)

Projectile diameter (km)

Total Os projectilea (moles)

Average Ir inventoryb (moles)

M(Os)impact (moles dissolved)

Soluble Os(%)c

Popigai Chicxulub Manicouagan

90 (a) 150 (a) 85 (a)

8 (b) 14 (c) n.a.

2.2  109 (d) 1.2  1010 (e) n.a.

1.5  108 (f) 1.5  109 (g) n.a.

0.7–1.5  108 (h) 0.4–2.3  109 (i) 2.8  108 (j)

3–7% 3–19% n.a.

Data Sources: (a) http://www.passc.net/EarthImpactDatabase/Diametersort.html; (b) Ivanov (2005); (c) Artemieva and Morgan (2009); (d) For Popigai, 611 ng/g Os (H chondrite: Kyte et al., 2011) is used rather than 843 ng/g Os (L chondrite: Tagle and Claeys, 2005) to minimize projectile Os inventory; (e) For Chicxulub, 614 ng/g Os (CM carbonaceous chondrite: Kyte, 1998; Goderis et al., 2013) is used rather than CV, CO or CR to minimize projectile Os inventory; (f) 5.7 ng Ir/cm2: geometric mean (n = 12; no uncertainty quoted) Liu et al. (2009); (g) 55 ± 3 ng Ir/cm2: geometric mean (n = 52) Donaldson and Hildebrand (2001). (h) Range from Table 5. (i) After Paquay et al. (2008a), using a range of (0.137–0.16) for Rpost-impact. See text for more discussion. (j) Calculated from Os isotope excursion reported by Sato et al. (2013). a Calculated from projectile diameter assuming a spherical geometry, projectile density of 2.6 g/cc and a projectile Os concentration (See Sato et al., 2013 supplemental material) for each projectile type as specified in the data sources below. b Average global Ir inventory in moles Ir is calculated by multiplying the global average Ir fluence by the surface area of the Earth. c Data column 5 divided by data column 3.

mixing between soluble, impact-derived Os with nearly chondritic 187Os/188Os and more radiogenic Os dissolved in the global ocean prior to the impact event. Importantly, this calculation also assumes that the total ocean inventory of dissolved Os has remained constant over time and is equal to that of the modern ocean. The multi-site data set characterizing the Os isotope excursion associated with the Popigai impact provides a means of directly comparing M(Os)impactor to Ir fluence from the same sites (Table 5). The resulting comparison provides empirical evidence that Os-based estimates of projectile size are more precise than those based on Ir data. The largest value of M(Os)impactor is a factor 2.5 times larger than the smallest M(Os)impactor estimate, while the ratio of maximum to minimum Ir fluences for these same sites is significantly larger (>12). This greater variability in Ir fluence compared to M(Os)impactor translates directly to variability associated with projectile size estimates (Table 6). M(Os)impactor is expected to be less variable than Ir fluence because the Os isotope composition of the open ocean, as recorded by marine sediments, is nearly homogenous at a give point time, while Ir fluence at a given site is governed by local patterns of sedimentation. Disagreement exists over the value of using excursions in the marine Os isotope record to estimate projectile size. Based on computational simulations of the Chicxulub impact event, Morgan (2008) argued that Os-based estimates of Chicxulub projectile size are gross underestimates (between 5 and 10-fold in projectile mass). Paquay et al. (2008b) responded by pointing out that the simulations cited by Morgan predict that roughly 80% of the Ir (and presumably Os) carried by the Chicxulub projectile remain within or near by the Chicxulub crater; a prediction that is not substantiated by available data. Note that for the Chicxulub impact, estimates of Rpost-impact vary and the lowest measured value (0.137: Meisel et al., 1995) closely approaches Rimpactor (0.13), leading to significant error amplification (Table 6). Specifically a 13% decrease in the value of Rpost-impact results in a nearly 4-fold increase in M(Os)impactor. Efforts to better constrain the 187Os/188Os of Rpost-impact are ongoing and will be reported separately.

Error amplification associated with estimates of M(Os)impactor for the Popigai impact is greatly diminished compared to that associated with Chicxulub (Table 5). Nevertheless, the sensitivity of M(Os)impactor to the value of Rpost-impact highlights the need for work establishing whether projectile-derived Os carried in insoluble phases has influenced our estimates of Rpost-impact for the Popigai impact. This is clearly important for the Manicouagan impact (Sato et al., 2013) where 187Os/188Os is indistinguishable from the chondrite value has been measured in a spherule-rich horizon. For Popigai, the fact that 3 of the 4 sites considered here have a very similar minimum measured 187 Os/188Os value (0.24–0.25) that is distinctly higher than chondrites suggests that Os inventory within Popigai impact horizon is dominated by Os scavenged from seawater. The apparent deficit in both Os and Ir that Morgan (2008) highlighted for Chicxulub is mirrored at Popigai, an impact event roughly 5 times smaller (Table 6). To make accurate estimates of projectile size, the fraction of projectile Os that actually dissolves in seawater “f(Os)soluble” must be known. In our original work, projectile size estimates were made assuming values of f(Os)soluble (0.3–1), using the equation below where “Mp(Os)” corresponds to the mass of the projectile, M(Os)impactor (as above) is the mass of projectile-derived Os that dissolves in seawater, and “C(Os)projectile” corresponds to projectile Os concentration. h i h i Mp ðOsÞ ¼ MðOsÞimpactor = fðOsÞsoluble  CðOsÞprojectile ð1Þ If we assume that estimates of projectile size based on models of crater formation (Ivanov, 2005: Artemieva and Morgan, 2009) are accurate, then estimates of f(Os)soluble for both the Popigai and Chicxulub impact events can be calculated (Table 6). The small tabulated values of f(Os)soluble we obtain imply that the 80–95% of impact-derived Os remains undissolved in seawater and this is a quantitative expression of the potential underestimate of projectile size. Note that the Os/Ir ratio of chondrites is 1.1 (Horan et al., 2003). Therefore the fact that estimated Ir inventories are essentially equal to M(Os)impactor for both Popigai

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and Chicxulub (Table 6, data columns 4 and 5) demonstrates that the magnitude of the deficit for both elements is are indistinguishable from one another for both Popigai and Chicxulub. Progress toward understanding the apparent deficit in measurable impact-derived Os and Ir might be achieved in several ways. First, if large inventories of Os and Ir can be found within, or close to, large impact craters, this would strongly support simulations of impact crater formation and the claim that Os-based estimates of projectile size are gross underestimates. Second, better constraining the oxygen fugacity of the expanding vapor plume produced by large impact events like Popigai would be useful in predicting whether or not formation of the volatile and soluble species, OsO4, occurs. If no OsO4 forms, then the very small values of f(Os)soluble implied by impact simulations are likely correct. Thermodynamic calculations designed to approximate conditions within the Chicxulub vapor plume indicate that conditions were sufficiently oxidizing to form the volatile species OsO4 (Ebel and Grossman, 2005), but to our knowledge no comparable calculations are available for Popigai or any other impacts. Third, there are opportunities to investigate other impact events. Recently Sato et al. (2013) characterized the signature of the Manicouagan impact, a crater whose diameter is very similar to that of the Popigai crater, in the marine Os isotope record. The value of M(Os)impactor is at least twice that we have estimated for Popigai but still much less than the total burden of Os that might be predicted for a crater nearly 100 km in diameter. Sato et al. (2013) emphasize the strong influence that the assumed Os concentration of the projectile has on projectile size estimates, highlighting the need to better constrain the type of impactor responsible for the Manicouagan crater as has been done for both Chicxulub and Popigai (See Table 6 notes). 5.3. The pre-impact decline in the marine Os record High temporal resolution Os isotope data from three sections reveal a pre-impact decline in bulk sediment 187 Os/188Os. Data from Sites 1090 and 1219 and the Massignano section show that the 187Os/188Os starts declining before the large positive Ir anomaly related to the Popigai impact (Figs. 2–4). Considering all three records we estimate that seawater 187Os/188Os decreased from 0.55 to 0.47 from the onset of the decline to immediately prior to the Popigai impact event. Based on comparison of 3He/4He in the Massignano section to the marine 187Os/188Os record, we argue that this decline is related to an episode of increased IDP flux that began approximately two hundred thousand years before the Popigai impact event (Farley et al., 1998; Farley, 2009). Elevated 3He/4He ratio is a sensitive indicator of the influence of interplanetary dust particles (IDPs) in marine sediments and has been used to recognize an episode of increased ET flux to the Earth in the late Eocene (Farley et al., 1998). Thus we interpret our data assuming that declining 187Os/188Os associated with elevated 3He/4He in the sediment record can be reasonably interpreted as the influence of IDPs on the marine Os isotope record.

Direct comparison of He and Os isotope variations from the Massignano section (Fig. 5) shows that the initial decline in 187Os/188Os coincides with a broad peak in 3 He/4He within C16n.1r. This feature is about 1 m below the 5.65 msl impact horizon in the Massignano section. The onset of this 187Os/188Os decline at Massignano is nearly synchronous with decreases in 187Os/188Os at other sites that also occur within magnetochron C16n.1r. At Site 1090 (Fig. 3), 187Os/188Os ratios begin to decline close to 294 meters composite depth (mcd) in parallel with gradual increases in Os and Ir concentrations. This is approximately 3 m below the impact horizon at 291.07 MCD. At Site 1219 (Fig. 4) it is difficult to resolve the exact depth where the pre-impact decline begins, but 187Os/188Os ratios become lower and distinctly more variable in the middle of C16n.1r at 182.75 revised meters composite depth (rmcd). This feature includes a brief excursion to low 187Os/188Os (0.34) that is associated with high Ir and Os concentrations at 182.46 rmcd, but it is distinct from the highest Ir concentrations at Site 1219 (1.7–0.6 ng/g Ir) that mark the impact horizon between 181.81 and 181.93 rmcd. In the age domain, 187Os/188Os ratios begin to decline at nearly the same time at all three sites; 36.07, 35.91 and 35.96 Ma at Massignano, Site 1090 and Site 1219, respectively. We conclude that the pre-impact decline in 187Os/188Os is more likely to be a primary depositional feature than a diagenetic feature (downward diffusion or physical mixing of impact-derived Os) because there is a better coherence between Os records in the time domain (see previous paragraph.), than in the depth domain. The pre-impact Os decline occurs over a 3 m depth interval at Site 1090, a 1–2 m interval at Massignano, and an approximately 1 m interval at Site 1219. For comparison, the length scale of downward mobility of impact-derived Os at the Cretaceous-Paleogene boundary in pelagic sequences is less than 60 cm. (See Fig. 7 of Ravizza and Vonderhaar, 2012.) Given the evidence that the pre-impact decline the in marine Os isotope record is a primary depositional feature that

Fig. 5. Bulk sediment 187Os/188Os ratios from the Massignano section from this study (small open squares) and Ravizza and Peucker-Ehrenbrink (2003) (filled large squares) and smoothed (5point running average) 3He/4He ratio (connected open circles: Farley et al., 1998) shown in stratigraphic height. Note that 187 Os/188Os is declining throughout C16n.1r, coincident with elevated 3He/4He and well before the Popigai impact horizon in C16n.1n. Magnetostratigraphy after Jovane et al. (2007).

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is concurrent with the previously recognized episode of elevated 3He flux at Massignano, we argue this feature of the marine Os isotope record is best interpreted as a consequence of elevated IDP flux in the late Eocene. 5.4. Estimates of the amount of soluble Os associated with the pre-impact decline in the marine Os isotope record A simple mass balance calculation allows estimation of the increase in the flux of soluble ET Os to the ocean that was associated with the pre-impact 187Os/188Os decline from 0.55 to 0.47. This approach assumes that the 187 Os/188Os variation we have measured in these various sediment sequences closely approximate the 187Os/188Os of contemporaneous seawater, and thus the calculations made are based upon our current understanding of the fluxes of soluble Os into and out of the global ocean. In order to make this calculation we construct three steady state marine Os budgets: one for the modern ocean, a second at the start of the pre-impact decline, and a third immediately prior to the Popigai impact event (Table 7). A two component mixing equation is used to approximate the marine Os isotope balance of the ocean corresponding to each steady state, Rsw ¼ Rr  f r þ Ru  ð1  f r Þ

ð2Þ

where R refers to 187Os/188Os and the subscript “sw” indicates seawater. The subscripts “r” and “u” indicate riverine and unradiogenic (combined mantle and background ET) sources of Os to seawater, respectively. The term fr refers to the mole fraction of the total 188Os flux (moles/yr) to the ocean that is supplied by rivers. f r ¼ Fr =ðFr þ Fu Þ

ð3Þ

The subscripts are defined as in Eq. (2) above and the term “F” is the flux of 188Os delivered to the ocean (moles/yr). The key assumptions in constructing each of the 3 steady states are as follows: (1) For the modern ocean, Fu can be estimated by difference using the riverine Os flux proposed by Sharma et al. (2007) and Ru = 0.13 and Rsw = 1.06. (2) For the onset of the pre-impact decline, Rr is revised downward relative to the modern value, using modern values of Fr and Fu to approximate Os fluxes to the ocean at this point in the late Eocene. This is done to account for the fact that our best estimate of Rsw = 0.55, about half that of the modern ocean. (3) Immediately prior to the Popigai impact, a new value of Fu is calculated assuming that riverine input (Fr) is unchanged and Rsw = 0.47. These calculations also

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assume that terrestrial Os inputs to the ocean were unchanged from the start of the pre-impact decline until the Popigai impact. The estimated flux of unradiogenic Os (Fu) to the ocean nearly doubles between onset of the pre-impact decline and the Popigai impact (Table 7). It is important to note that this analysis implicitly assumes that this additional flux of ET Os is soluble in seawater. The integrated flux of excess unradiogenic Os over the approximately 200 kyr between the start of the pre-impact decline in 187Os/188Os and the Popigai impact can be estimated by assuming that the flux of excess unradiogenic Os to the ocean increases linearly from zero at the start of the pre-impact decline to 53 moles 188 Os/yr just prior to the impact (Table 7). The integrated total influx of excess ET Os dissolved in the ocean over 200 kyr is 4  107 moles Os, which is 1/4 to 2/3 of the amount of Os instantaneously added to the seawater Os inventory by the Popigai impact event (see range of M(Os)impactor for Popigai in Table 6). Thus, although the episode of increased ET flux to Earth that is marked by elevated 3He flux also affected the marine Os isotope balance, the integrated magnitude of the total soluble Os flux to the ocean was smaller than that supplied by the Popigai impact event. The calculations above address only the soluble component of the ET Os flux, but there is also a particulate component of ET Os flux that needs to be considered. The background flux of particulate ET matter has been shown to be important in sites with low total Os burial fluxes (Esser and Turekian, 1988; Ravizza and McMurtry, 1993; Peucker-Ehrenbrink, 1996). The contribution of particulate ET Os to marine sediments can be estimated using the following equation: (Particulate ET Os) = (Os chondrites)  (3He measured/3He IDP/0.005) (Marcantonio et al., 1999; Dalai and Ravizza, 2006) where average chondritic Os concentration is 492 ng/g (Tagle and Berlin, 2008), the measured 3He is from the Massignano section (Farley et al., 1998), and 3He concentration in IDP is 1.9  105 cc STP/g (Nier and Schlutter, 1992). The factor 0.005 is the estimated fraction of ET He (0.5% of total mass) that survives upon atmospheric entry (Marcantonio et al., 1995; Farley et al., 1997). This approach makes the assumption that the 3He/Os ratio of late Eocene IDPs can be approximated by that estimated for modern IDPs; an assumption that may not be valid because the ET flux of He and Os is carried primarily by IDPs of different sizes. ET 3He is carried mainly by those IDPs too small to heat significantly upon atmospheric entry; IDPs < 35 lm which

Table 7 Marine Os budgets constructed to estimate the soluble Os delivered by episode of increased IDP flux for the 200,000 years before the Popigai impact. Riverine Os

a b

Os/188Os

Seawater

187

188

187

187

170 170 170

1.38 0.69b 0.69

58b 58 111b

0.13 0.13 0.13

1.06 0.55 0.47

Os moles/yr

Modern Oceana Pre-impact decline (start) Immediately (end)

Unradiogenic Os

188

Os moles/yr

River flux and 187Os/188Os from Sharma et al. (2007). Calculated using Eqs. (2) and (3) and other parameters given in the row in which the entry occurs.

Os/188Os

Os/188Os

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Popigai Impact

0.7

0.6

0.5

188

Os

Eocene-Oligocene Boundary

Massignano as the influence of unradiogenic Os carried by mineral detritus. This interpretation is consistent with both the relatively poor analytical reproducibility characteristic of Massignano samples (Section 4.1), and our inference that Massignano samples may carry discrete Ir-rich detrital particles (Section 5.1). The ODP 1090 and ODP 1218/1219 record display less short term variability in 187Os/188Os than Massignano, allowing an offset between these two records immediately following the Popigai impact to be confidently identified. Recent work indicating that the 187 Os/188Os of modern seawater may not be homogeneous raises the possibility that the higher 187Os/188Os in the Equatorial Pacific (ODP 1218/1219) relative to the South Atlantic (OPD 1090) may reflect regional contrasts in late Eocene seawater at about 35.5 Ma (Fig. 6). Although the available data do not allow us to draw any robust conclusions regarding the homogeneity of 187Os/188Os in late Eocene ocean, we highlight this as a topic that requires further investigation. In spite of differences between records discussed above, all three records show coherent variation between sites; in particular, two pronounced minima in 187Os/188Os are readily apparent. The Os isotope excursion centered at 35.8 Ma is very brief compared to the excursion of similar amplitude at centered at 34.5 Ma. As detailed above (Sections 5.2 and 5.3), the feature at 35.8 Ma results from the combined influence of the episode of increased IDP flux, referred to as the pre-impact decline in the previous sections, and the Popigai impact event. The source of unradiogenic Os responsible for the more prominent 34.5 Ma feature remains uncertain. Previous studies suggest either an increased influx of extraterrestrial material (Ravizza and Peucker-Ehrenbrink, 2003; Dalai et al., 2006) or accelerated weathering of ophiolitic sequences (Pegram and Turekian, 1999; Reusch, 2011) may have led to lowered 187 Os/188Os at 34.5 Ma. The former explanation links the

Os/

represent less than 0.5% of the total ET mass flux (Farley et al., 1997; Mukhopadhyay and Farley, 2006). In contrast, the ET Os flux is carried by the entire range of ET particle sizes and is dominated by the 100–200 lm size range. Therefore, conditions that modify the size distribution of IDPs reaching the Earth are also likely to modify the 3 He/Os ratio of the bulk ET flux to Earth. Sediment inventories of particulate ET Os estimated using the approach described in the preceding paragraph are too low and too variable to persistently influence the bulk sediment Os budget during the studied interval based upon Os burial flux estimates at Massignano. The calculated contribution of particulate Os in the Massignano samples varies from 0.2 to 3 pg/g, corresponding to a range of 0.2–14% of the total Os inventory of the sediment. The particulate Os burial flux is then calculated from the bulk density of 2.5 g/cm3 (Farley et al., 1998) and an average sedimentation rate over the studied interval (1.0 cm/ kyr). The particulate Os burial flux varies from 0.39 to 5.1 pg/cm2/kyr. This calculated particulate ET flux is within error of Late Cretaceous values (1.4 ± 0.6 pg/cm2/kyr: Ravizza, 2007), early Pleistocene values (1.21 ± 0.47 pg/ cm2/kyr: and Dalai and Ravizza, 2010) and of recent sediment (3.5 pg/cm2/kyr: Esser and Turekian, 1988; Ravizza and McMurtry, 1993). However, individual particles of ET matter may contribute to poor reproducibility in some instances (See Section 4 above). We infer that much of the Os carried by the late Eocene IDP pulse must have been soluble in seawater because our estimates of particulate ET Os can account for only a small and highly variable fraction of the total Os observed in the sediment. This conclusion has potentially important implications for the modern marine Os budget. Early efforts to construct an Os isotope balance for the modern ocean argued that alteration of mafic and ultramafic rocks, rather than dissolution of ET particles, was the primary source of unradiogenic Os to seawater (Levasseur et al., 1999; Peucker-Ehrenbrink and Ravizza, 2000). However, analyses of Os in seafloor hydrothermal fluids reveal generally low Os concentrations, necessitating a reassessment of the marine Os budget that allows for the possibility that dissolution of particulate ET Os in seawater could be the major source of unradiogenic Os to seawater (Sharma et al., 2007). Results from this study of the late Eocene marine Os isotope record support the idea that ET matter provides a significant fraction of the total flux of soluble Os to the ocean. However, caution is warranted in extrapolating from the late Eocene, a time of elevated ET flux to the Earth, to the modern.

187

254

0.4

ODP 1218/1219

0.3

ODP 1090

Massignano

5.5. A composite marine Os isotope record of the late Eocene Examining the 187Os/188Os records from the three sites studied here on a common age scale (Fig. 6) reveals considerable scatter, as well as one interval where there are systematic off sets between the records. The Massignano record displays the most variation over short time intervals among the three sites and, in general, measured 187Os/188Os ratios are lower than either ODP 1090 or ODP 1218/1219 at a given time. We interpret the relatively low 187Os/188Os at

0.2 36.5

36

35.5

35

34.5

34

33.5

33

32.5

Age (Ma)

Fig. 6. Composite late Eocene marine Os isotope record based on Massignano (this study, Ravizza and Peucker-Ehrenbrink, 2003), ODP 1218/1219 (this study, Dalai et al., 2006 and Paquay et al., 2008a) and ODP 1090 (this study, Paquay et al., 2008a). Age assignments are based on magnetostratigraphic datums given in Table 1.

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negative excursion in 187Os/188Os at 34.5 Ma to the episode of increased IDP flux identified by the 3He/4He maximum at Massignano (Fig. 5) by invoking Poyinting–Robertson (P– R) drag to fractionate smaller He-rich particles (<20 lm) from the larger particles (100–200 lm) that dominate the ET flux of Os and Ir to the Earth. The P–R drag phenomenon is required to explain the roughly one million year delay between the 3He/4He peak and the 187Os/188Os minimum. However, results reported here strongly suggest that an increased influx of soluble ET Os accompanied increased 3 He flux just prior to the Popigai impact. This argues against the suggestion of Ravizza and Peucker-Ehrenbrink (2003) that P–R drag caused a lag between the time of maximum ET He and ET Os flux in the late Eocene. Instead we prefer to attribute the prominent 187Os/188Os minimum that occurs a few hundred thousand years before the Eocene–Oligocene transition, to changes in terrestrial Os inputs to the ocean. The results of this study therefore lend support to a proposed link between the prominent 187Os/188Os minimum at approximately 34.5 Ma (Fig. 5), and global changes in weathering proposed by Reusch (2011) as a contributing factor to the Eocene/Oligocene transition from a greenhouse to an icehouse world.

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the result of IDP dissolution as suggested by Ravizza and Peucker-Ehrenbrink (2003). Instead, globally significant changes in terrestrial weathering fluxes (Reusch, 2011) are likely responsible for this longer and larger perturbation of the marine Os isotope cycle. In addition, this finding hints that dissolution of Os from the background flux of ET matter may exert a significant influence on the 187 Os/188Os of seawater at times other than the late Eocene. ACKNOWLEDGMENTS FSP thanks Steven Bohaty for suggesting studying ODP Site 1218, Bernhard Diekmann, Jennifer Latimer, Annette Olivarez Lyle and Thomas Westerhold for sharing data from ODP Sites 1090, 1218 and 1219. This research used samples and data provided by the Integrated Ocean Drilling Program (IODP); in particular we thank Walter Hale (Bremen repository) for his patience and excellent work with the numerous sample requests. Discussion with Steven Goderis and James Hetfield were useful. We thank Philippe Claeys, Frank Kyte, Sandro Montanari, one anonymous reviewer and associate editor Christian Koeberl for thoughtful comments on various versions of this manuscript, which helped us to improve the final product. Denys Vonderhaar provided essential laboratory support for this work. NSF awards, EAR0843930 and OCE1061061, to GR supported this research.

6. CONCLUSIONS At Massignano, ODP 1090 and ODP 1219/1218 there is a single prominent excursion to low 187Os/188Os is that is correlated with elevated Ir concentrations. This excursion occurs within magnetochron C16n.1n at all study sites, indicating that this feature in the marine Os isotope record is a consequence of the Popigai impact event. The absence of similar Os isotope excursions at 6.2 and 10.3 msl at Massignano argue against previous claims that these horizons record additional impact events. Comparison of Os and Ir data from these sites demonstrates that Os-based estimates of projectile size are far less variable than those derived from Ir fluence measurements. The new Os isotope data substantiate our earlier conclusion (Paquay et al., 2008a) that only a small fraction (3–7%) of the total quantity of Os implied by models of Popigai crater formation dissolved in seawater. A compilation of Ir fluence estimates (Liu et al., 2009) indicates a similar deficit in measurable Ir inventory. The fractional deficit of measurable impact-derived Os and Ir associated with the Popigai impact is similar to that established for the Chicxulub impact. Finding evidence of this heretofore undetected projectile-derived Os and Ir would represent an important contribution to validating computational models of impact crater formation. Careful comparison of 3He/4He (Farley et al., 1998) and 187 Os/188Os records from the Massignano section shows that a decline in 187Os/188Os starts begins at nearly the same time as a peak in 3He flux, but prior to the Popigai impact. A similar pre-impact decline in 187Os/188Os is also present at ODP 1090 and ODP 1218/1219. These results suggest that the episode of increased IDP flux recorded by He isotopes delivered a sufficiently large quantity of soluble Os to the ocean to affect seawater 187Os/188Os. An important consequence of this interpretation is that a second prominent excursion to low 187Os/188Os at 34.5 Ma is unlikely to be

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