Supplemental Materials: Impact of Dissolution, Bioturbation and Winnowing on the Sedimentary Record of the Paleocene-Eocene Thermal Maximum
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- 2. Products of Dissolution
- 3. Interpreting bioturbation and winnowing from single specimen and bulk records
| Supplemental Materials: Impact of Dissolution, Bioturbation and Winnowing on the Sedimentary Record of the Paleocene-Eocene Thermal Maximum
1. Quantifying dissolution during the PETM In Supplemental Figure 1 we plot percent dissolution versus paleodepth for all PETM sites with available CaCO3 data as compiled by Panchuk (2007), with the addition of the new data from Site 690 and data from Site 401 from Cecily Palike (pers. comm., 2013) (Table 1). We exclude Site 865 from the plot as the PETM section is clearly incomplete (e.g., Bralower et al., 1995). These data suggest that Shatsky Rise Sites 1209-1212 were apparently characterized by less intense dissolution than other sites at comparable paleodepths with the exception of Sites 690 and 401 (N. Atlantic). We propose that the impact of dissolution on carbonate at Site 1209 has been diminished by bioturbation as discussed in Section 4.2. However, changes in the fluxes of carbonate may also have played a role.  Supplemental Materials Figure 1. Percent dissolution determined using the technique of Broecker (1995). Red numbers are Pacific Sites, blue are sites in the Atlantic, Caribbean and Southern Ocean. See Table 1 for data.
Table 1. Dissolution percent calculated from pre (high) and peak (low) CaCO3 contents after Broecker (1995). Sites 690, 1209 and 1262 and all PETM sections are characterized by predominantly biogenic sediment sources (largely nannoplankton and planktic foraminifera) with minor contributions from aeolian and clastic sources. In all likelihood, fluxes of carbonate and non-carbonate particles changed during the PETM with variation in plankton production, weathering and aeolian transport (e.g., Robert and Kennett, 1994; Ravizza et al., 2001; Bralower, 2002; Kelly, 2002; Gibbs et al., 2006; Petrizzo, 2007). However, there is no way to determine exactly the changes in these fluxes, nor can we estimate the loss of highly fragile plankton species that are only preserved in clay-rich, hemi-pelagic deposits along continental margins (e.g., Bown et al., 2008). The dissolution estimates (Figure 6 and Supplemental Figure 1) are based on assuming that the fluxes from the different sources remained constant. At the three study sites the amount of dissolution estimated from the change in percent CaCO3 (Broecker, 1995, 2009; Stap et al., 2009; Stap, 2010) at the peak of the PETM is over 60%, over 70%, and 100% of the carbonate at Sites 690, 1209, and 1262, respectively (Figures 6a,b). At Site 1262, complete dissolution is clear from the absence of CaCO3 in the interval corresponding to the peak of the PETM. At Site 690, impoverished nannofossil preservation supports the notion of a significant amount of dissolution, quite possibly approaching 60%. However, nannofossil preservation at Site 1209 does not reveal dissolution to the extent of 70% of the sediment (Plate 1). As discussed bioturbation has likely masked the extent of dissolution at this site. In addition, the assumption of constant production of carbonate has led to an overestimate of the calculated levels of dissolution at Sites 690 and 1209. Plankton assemblages suggest oligotrophic conditions in open-ocean locations during the PETM (Bralower, 2002; Gibbs et al., 2006), and carbonate production rates at these sites are likely to have decreased modestly during the initial stages of the PETM (e.g., Kelly et al., 2012). Gibbs et al. (2010) estimated a carbonate production rate decrease of ~15-20% at Site 690 and ~5% at Site 1209. However, GENIE model simulations suggest that the biological pump was less efficient at Site 690 than at Site 1209 where the water column was stratified in the early part of the PETM (Schneider et al., 2013). Thus the decrease in carbonate supply might have been more significant at Site 1209. At Site 690, clay accumulation may have accelerated as a result of enhanced continental weathering on the nearby Antarctic continent (Robert and Kennett, 1994). 2. Products of Dissolution When sediment is chemically eroded, a significant amount of the dissolved carbonate is released into pore waters. At Site 1209, the dissolved carbonate appears to have reprecipitated as long 5-15 µm calcite blades (Colosimo et al., 2006) (Supplemental Plate 1). Large blocks of reprecipitated carbonate hosting whole foraminiferal shells and molds of foraminifera have been observed at Site 865, another site where the brevity of the PETM interval is likely partially a result of burndown (Supplemental Plate 1, see similar objects in Kozdon et al., 2013). At Site 1262, small rhombs of dolomite or calcite are observed in samples with low or no carbonate; these rhombs are possibly precipitates of dissolved carbonate. At Site 690 it is impossible to distinguish between micarb and clay in electron micrographs, but at >50% CaCO3, some of the significant volume of fine grained material (Plate 1) must be micarb derived from reprecipitation of the dissolved CaCO3. Supplemental Materials Plate 1. Electron micrographs showing nannofossil preservation across the base of the CIE at Site 690. The micrographs can distinguish the abundance and preservation of coccolith carbonate, but cannot generally distinguish between clay and micarb. a. 170.53-170.65 mbsf; b. 170.67-170.73 mbsf; c.170.74-170.79 mbsf; d. 170.80-170.82 mbsf. Samples below 170.80 show moderate dissolution with samples dominated by coccoliths and pieces of coccoliths and a lower volume of clay/micarb. Between 170.80 and 170.75 mbsf there is a substantial decline in preservation with fewer coccoliths with more etched rims and more clay/micarb. Samples are poorly preserved up to 170.70 mbsf. Coccolith preservation is very poor from 170.69 to 170.65 mbsf with few whole coccoliths and a lot of clay/micarb. Samples from 170.64 mbsf and above have an increasing number of whole coccoliths with a lower volume of clay/micarb. Onset of the bulk CIE is at 170.70 mbsf (Figure 2). Supplemental Materials Plate 2. Electron micrographs showing nannofossil preservation across the base and lower part of the CIE at Site 1209 (Section 22H-1). a. 196.47-196.51 mbsf; b. 196.46-196.44 mbsf; c. 196.28 and 196.42 mbsf. Samples are dominated by nannofossils. For most of the section, fossil preservation is moderate. There is a slight decline in preservation between 196.46 and 196.47 mbsf. Lithologic boundary and CIE are at 196.45 mbsf (Figure 3). Supplemental Materials Plate 3. Morphology of calcite likely derived from reprecipitation of calcite dissolved in the PETM. a. and b. Calcite blades on planktonic foraminifera from Site 1209. c. and d. Blocks of reprecipitated calcite hosting foraminifera from Site 865. Scale bars in c. and d. are 100 µm. 3. Interpreting bioturbation and winnowing from single specimen and bulk records In Section 4.3 we discuss evidence for bioturbation and winnowing, including single specimen foraminiferal and bulk stable isotopes. In Supplemental Materials Figure 2a we show single specimen planktonic isotope and bulk isotope records from Site 690 (data from Thomas et al., 2002). As discussed in the text, the single specimen data show 10 specimens that are clearly out of place, two specimens of Subbotina that have been mixed upwards and eight specimens of Acarinina that have been mixed downward. Any of these specimens could be the result of bioturbation and winnowing. Bioturbation is most likely to mix foraminiferal specimens downward while winnowing has the potential to mix specimens upwards, thus the Subbotina specimens are more likely mixed via winnowing and the specimens of Acarinina mixed via bioturbation. It is unlikely that bioturbation has transported these specimens by 20 cm. This mixing is more likely the result of downhole contamination during drilling. 
Supplemental Materials Figure 2a. Distribution of single specimen and bulk carbonate carbon isotope values at Site 690 (data from Thomas et al. 2002). All samples are from the archive half of the core. Circles designate samples out of place potentially as a result of bioturbation or winnowing. Lines indicate levels of samples with grain size distributions potentially indicative of winnowing. 
Supplemental Materials Figure 2b. Distribution of single specimen oxygen isotope values at Site 690 (data from Thomas et al. 2002). All samples are from the archive half of the core. Circles designate samples out of place potentially as a result of bioturbation or winnowing. Lines indicate levels of samples with grain size distributions potentially indicative of winnowing. Grain size data allow us to identify samples that have distributions potentially altered by winnowing (Supplemental Figure 3). We indicate these samples in Supplemental Materials Figure 2. The lack of coincidence between samples with isotope values that are out of place and those that have grain size distributions consistent with winnowing suggests that winnowing is not the cause of displacement. Possibly, the two specimens of Subbotina have been mixed downwards by drilling. From the single specimen isotope data it is possible that a number of specimens that are not clearly out of place have been mixed by bioturbation and winnowing, just not to the extent that would distinguish them from surrounding specimens. This is particularly the case close to the carbon isotope excursion. 
Supplemental Materials Figure 3. Logarithmic grain size profiles for samples from across the base of the PETM at the three study sites. 10% scale bar is volume abundance of the sediment fraction. The data show unimodal and relatively homogenous distributions at Sites 1209 and 1262; both sites show a shift towards finer grain size in samples from the earliest part of the PETM (196.45-196.42 mbsf at Site 1209; 140.1-139.9 mcd at 1262). Samples at Site 690 generally show a more bimodal grain size distribution than the other two sites, with most samples possessing a very minor but still distinguishable peaks likely representing nannoplankton and foraminifera. Some samples (including 170.36, 170.45, 170.53, 170.69, 170.725, 170.755, 170.845 mbsf) show a more pronounced foraminiferal peak. The impact of bioturbation can clearly be observed in bulk carbonate isotope data from Site 690. We plot bulk carbonate isotopes data from the two halves of Section 19-1 in Supplemental Materials Figure 3. The plot shows largely comparable values between 170.4 and 170.8 meters below sea floor with the exception of the interval between 170.62 and 170.69 mbsf where a burrow has clearly mixed the bulk carbonate isotope values in the archive half with respect to the working half. 
Supplemental Materials Figure 4. Bulk carbonate carbon isotope values of the archive half of the PETM (Section, 690B-19H-1) (from Thomas et al., 2002) plotted against values from the working half (from Bains et al., 1999).
Supplemental Materials Table 2. Level of significant changes associated with the PETM at the three study sites (See Figure 2) References Bains, S., Corfield, R.M., Norris, R.D., 1999. Mechanisms of climate warming at the end of the Paleocene: Science, v. 285, p. 724–727. Bown, P.R., Dunkley Jones, T., Lees, J.A., Randell, R.D., Mizzi, J.A., Pearson, P.N., Coxall, H.K., Young, J.R., Nicholas, C.J., Karega, A., Singano, J., Wade, B.S., 2008. Paleogene calcareous microfossil Konservat-Lagerstätte from the Kilwa Group of coastal Tanzania Geological Society of America Bulletin, 120, 3-12, doi:10.1130/B26261.1 Bralower, T. J., 2002. Evidence for surface water oligotrophy during the late Paleocene thermal maximum: nannofossil assemblage data from Ocean Drilling Program Site 690, Maud Rise, Weddell Sea, Paleoceanography, 17, doi 10.1029/2001PA000662. Bralower, T. J., Zachos, J.C., Thomas, E., Parrow, M., Paull, C.K., Kelly, D.C., Premoli Silva, I., Sliter, W.V., Lohmann, K.C., 1995. Late Paleocene to Eocene paleoceanography of the equatorial Pacific Ocean: stable isotopes recorded at ODP Site 865, Allison Guyot: Paleoceanography, 10, 841-865.
Broecker, W., 2009, Wally’s Quest to Understand the Ocean’s CaCO3 Cycle, Annual Review Marine Science, 1:1–18, doi:10.1146/annurev.marine.010908.163936 Colosimo, A., Bralower, T.J., Zachos, J. 2006. Evidence for lysocline shoaling at the Paleocene/Eocene Thermal Maximum on Shatsky Rise, Northwest Pacific. In Bralower, T. J., I. Premoli-Silva, M. J. Malone, et al. Proc. ODP Sci. Res., 198, 1-36.
Gibbs, S.J., Stoll, H., Bown, P.R., Bralower, T.J., 2010. Ocean acidification and surface water carbonate production across the Paleocene-Eocene Thermal Maximum, Earth and Planetary Science Letters 295, 583-592. Kelly, D. C., 2002. Response of Antarctic (ODP Site 690) planktonic foraminifera to the Paleocene-Eocene thermal maximum: faunal evidence for ocean/climate change. Paleoceanography 17, doi 10.1029/2002PA000761. Kelly, D. C., Nielsen, T. M. J., Schellenberg, S. A., 2012. Carbonate saturation dynamics during the Paleocene-Eocene thermal maximum: bathyal constraints from ODP Sites 689 and 690 in the Weddell Sea (South Atlantic). Marine Geology, doi: 10.1016/j.margeo.2012.02.003. Kozdon, R., Kelly, D.C., Kitajima, K., Strickland, A., Fournelle, J. H., Valley, J.W., 2013. In situ δ18O and Mg/Ca analyses of diagenetic and planktic foraminiferal calcite preserved in a deep-sea record of the Paleocene-Eocene thermal maximum. Paleoceanography, DOI: 10.1002/palo.20048. Panchuk, 2007. Investigating the Paleocene/Eocene carbon cycle perturbation: An Earth System model approach. Unpublished PhD thesis, The Pennsylvania State University, 83pp. Petrizzo, M.R. 2007. The onset of the Paleocene–Eocene Thermal Maximum (PETM) at Sites 1209 and 1210 (Shatsky Rise, Pacific Ocean) as recorded by planktonic foraminifera. Marine Micropaleontology (2007), doi:10.1016/j.marmicro.2006.11.007 Ravizza, G., Norris, R. N., Blusztajn, J., Aubry, M.P., 2001. An osmium isotope excursion associated with the late Paleocene thermal maximum: Evidence of intensified chemical weathering. Paleoceanography, 16, 155-163. Robert, C. Kennett, J.P., 1994. Antarctic subtropical humid episode at the Paleocene-Eocene boundary: Clay-mineral evidence. Geology, 22: 211-214. Schneider, L. J., Bralower, T.J., Kump, L.R., Patzkowsky, M.E., 2013. Calcareous nannoplankton ecology and community change across the Paleocene-Eocene Thermal Maximum, Paleobiology, 39(4), 2013, pp. 628–647 DOI: 10.1666/12050. Stap, L., Sluijs, A., Thomas, E., Lourens, L., 2009. Patterns and magnitude of deep sea carbonate dissolution during Eocene Thermal Maximum 2 and H2, Walvis Ridge, southeastern Atlantic Ocean. Paleoceanography, 24, PA1211, doi:10.1029/2008PA001655. Stap, H.L., 2010. Palaeoceanographic and biotic response during early Eocene extreme global warming events, Geologica Ultraiectina, 328. Thomas, D. J., Zachos, J. C., Bralower, T. J., Thomas, E., Bohaty, S., 2002. Warming the fuel for the fire: evidence for the thermal dissociation of methane hydrate during the Paleocene-Eocene thermal maximum. Geology 30, 1067-1070. Download 139.16 Kb. Share with your friends:
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