cryogenian glaciation and the onset of carbon-isotope decoupling … · 2013. 7. 30. · values...

13. J. Helbert, A. Maturilli, N. Mueller, in VenusGeochemistry: Progress, Prospects, and New Missions(Lunar and Planetary Institute, Houston, TX, 2009), abstr.2010.

14. J. Helbert, A. Maturilli, Earth Planet. Sci. Lett. 285, 347(2009).

15. Coronae are circular volcano-tectonic features that areunique to Venus and have an average diameter of~250 km (5). They are defined by their circular and oftenradial fractures and always produce some form of volcanism.

16. G. E. McGill, S. J. Steenstrup, C. Barton, P. G. Ford,Geophys. Res. Lett. 8, 737 (1981).

17. R. J. Phillips, M. C. Malin, Annu. Rev. Earth Planet. Sci.12, 411 (1984).

18. E. R. Stofan, S. E. Smrekar, D. L. Bindschadler, D. Senske,J. Geophys. Res. 23, 317 (1995).

19. Estimated elastic thickness values at Themis Regio aretypically 10 to 20 km (20, 21). The authors of (22)conducted a global admittance study and found values ofelastic thickness of 0 to 50 km at both Dione andThemis Regiones. Their analysis also shows regions oflarge elastic thickness, up to 100 km, in the northernportion of Dione Regio covering Ushas Mons. Estimates ofaverage apparent depth of compensation (ADC) for Dioneand Themis Regiones are 130 km and 100 km (21),respectively. The elastic thickness at Imdr Regio cannotbe reliably estimated due to the low resolution of thegravity field in that region (53). Stofan et al. (18)estimated an ADC of 260 km, which is consistent with adeep plume.

20. M. Simons, S. C. Solomon, B. H. Hager, Geophys. J. Int.13, 24 (1997).

21. S. E. Smrekar, E. R. Stofan, Icarus 139, 100 (1999).22. F. S. Anderson, S. E. Smrekar, J. Geophys. Res. Planets

111, E08006 (2006).23. S. T. Keddie, J. W. Head, J. Geophys. Res. 101, 11,729

(1995).24. E. R. Stofan, J. E. Guest, A. W. Brian, Mapping of V-28

and V-53, T. K. P. Gregg, K. L. Tanaka, R. S. Saunders,

Eds. (U.S. Geological Society Open-File Report 2005-1271, Abstracts of the Annual Meeting of PlanetaryGeologic Mappers, Washington, DC, 2005), pp. 20–21.

25. E. R. Stofan, S. E. Smrekar, J. Helbert, P. Martin,N. Mueller, Lunar Planet. Sci. XXXIX, abstr. 1033(2009).

26. B. D. Campbell et al., J. Geophys. Res. 97, 16,249(1992).

27. R. J. Phillips, N. R. Izenberg, Geophys. Res. Lett. 22, 1617(1995).

28. C. M. Pieters et al., Science 234, 1379 (1986).29. B. Fegley Jr., R. G. Prinn, Nature 337, 55 (1989).30. A. H. Treiman, C. C. Allen, Lunar Planet. Sci. Conf. XXV,

1415 (1994).31. M. I. Zolotov, V. P. Volkov, in Venus Geology,

Geochemistry, Geophysics—Research Results from theUSSR (Univ. of Arizona Press, Tucson, AZ, 1992),pp. 177–199.

32. B. Fegley, A. H. Treiman, V. L. Sharpton, Proc. LunarPlanet. Sci. 22, 3 (1992).

33. B. Fegley, K. Lodders, A. H. Treiman, G. Klingelhöfer,Icarus 115, 159 (1995a).

34. B. Fegley et al., Icarus 118, 373 (1995b).35. A. M. Baldridge, S. J. Hook, C. I. Grove, G. Rivera,

Remote Sens. Environ. 113, 711 (2009).36. L. T. Elkins-Tanton et al., Contrib. Minerol. Petrol. 153,

191 (2007).37. S. A. Gibson, R. N. Thompson, A. P. Dickin, Earth Planet.

Sci. Lett. 174, 355 (2000).38. L. S. Glaze, J. Geophys. Res. 104, 18,899 (1999).39. New Scientist 23, 52 (2009) (www.newscientist.com/

article/dn17534).40. V. S. Meadows, D. Crisp, J. Geophys. Res. 101, 4595

(1996).41. B. Fegley, Icarus 128, 474 (1997).42. E. R. Stofan, A. W. Brian, J. E. Guest, Icarus 173, 312

(2005).43. P. R. Hooper, in Large Igneous Provinces: Continental,

Oceanic and Planetary Flood Volcanism, J. J. Mahoney,

M. F. Coffin, Eds. (Monograph 100, American GeophysicalUnion, Washington, DC, 1997), pp. 1–28.

44. M. A. Bullock, D. H. Grinspoon, Icarus 150, 19 (2001).45. R. G. Strom, G. G. Schaber, D. D. Dawson, J. Geophys.

Res. 99, 10,899 (1994).46. K. M. Roberts, J. E. Guest, J. W. Head, M. G. Lancaster,

J. Geophys. Res. 97, 15,991 (1992).47. C. R. K. Kilburn, in Encyclopedia of Volcanoes,

H. Sigurdsson, Ed. (Academic Press, San Diego, CA,2000), pp. 291–306.

48. We have rounded these numbers in recognition that theage estimates have higher uncertainties than the volumeestimates.

49. G. Choblet, E. M. Parmentier, Phys. Earth Planet. Inter.173, 290 (2009).

50. M. E. Davies et al., Celestial Mech. 39, 103 (1986).51. P. K. Seidelmann et al., Celestial Mech. Dyn. Astron. 82,

83 (2002).52. M. E. Davies et al., J. Geophys. Res. 97, 13,141

(1992).53. A. S. Konopliv, W. S. Banerdt, W. L. Sjogren, Icarus 139,

3 (1999).54. G. L. Hashimoto, T. Imamura, Icarus 154, 239 (2001).55. This research was carried out in part at the Jet Propulsion

Laboratory, California Institute of Technology, and wassponsored by the Planetary Geology and GeophysicsProgram and NASA. We gratefully acknowledge the workof the entire Venus Express and VIRTIS teams. We thankthe European Space Agency, Agenzia Spaziale Italiana,Centre National des Etudes Spatiales, CNRS/InstitutNational des Sciences de l’Univers, and the othernational space agencies that have supported thisresearch. VIRTIS is led by INAF-IASF, Rome, Italy, andLESIA, Observatoire de Paris, France.

7 January 2010; accepted 25 March 2010Published online 8 April 2010;10.1126/science.1186785Include this information when citing this paper.

Cryogenian Glaciation and the Onsetof Carbon-Isotope DecouplingNicholas L. Swanson-Hysell,1 Catherine V. Rose,1 Claire C. Calmet,1 Galen P. Halverson,2*Matthew T. Hurtgen,3 Adam C. Maloof1†

Global carbon cycle perturbations throughout Earth history are frequently linked to changingpaleogeography, glaciation, ocean oxygenation, and biological innovation. A pronouncedcarbonate carbon-isotope excursion during the Ediacaran Period (635 to 542 million yearsago), accompanied by invariant or decoupled organic carbon-isotope values, has been explainedwith a model that relies on a large oceanic reservoir of organic carbon. We present carbonate andorganic matter carbon-isotope data that demonstrate no decoupling from approximately 820 to760 million years ago and complete decoupling between the Sturtian and Marinoan glacialevents of the Cryogenian Period (approximately 720 to 635 million years ago). Growth of theorganic carbon pool may be related to iron-rich and sulfate-poor deep-ocean conditions facilitatedby an increase in the Fe:S ratio of the riverine flux after Sturtian glacial removal of a long-livedcontinental regolith.

Throughout most of the Phanerozoic Eon[542 million years ago (Ma) to present],paired records of carbonate carbon (d13Ccarb)

and coeval bulk organic carbon (d13Corg) iso-topes are consistent with a model in which theorganic carbon in marine sediments is derivedand fractionated from contemporaneous dissolvedinorganic carbon (DIC). In contrast, d13Ccarb andd13Corg records from Ediacaran (635 to 542 Ma)carbonate successions (1–3) show relatively in-

variant d13Corg during large changes to d13Ccarb

across the ~580 million-year-old Shuram-Wonokaanomaly (Fig. 1 and fig. S1). This behavior hasbeen used to develop and support a model forthe Neoproterozoic (1000 to 542 Ma) carboncycle in which invariant d13Corg values result froma very large oceanic reservoir of 13C-depleted dis-solved organic carbon (DOC) and particulateorganic carbon (POC) (or, alternatively, sourcedfrom a large sedimentary reservoir) that over-

whelms the signal from primary biomass frac-tionated from contemporaneous DIC (4). Weconsider the large oceanic reservoir model and,as in (2), use the term DOC to collectively referto organic carbon that is truly dissolved as wellas suspended colloidal organic carbon and finePOC. The buildup and maintenance of a largeDOC pool implies low Corg remineralization—perhaps associated with low oxygen and sulfatelevels—but high nutrient liberation efficiency.In such an ocean, the d13C of the DIC pool issensitive to inputs (via remineralization) from the13C-depleted DOC pool that can drive negativeexcursions. The end of the invariance in thed13Corg record in the latter stages of the Shuram-Wonoka anomaly has been interpreted as thedemise of the large DOC pool (2, 3).

Stratigraphically constrained coupled recordsof d13Ccarb–d

13Corg at sufficient detail to test thiscarbon cycle model have been available onlyfrom carbonates of Ediacaran age (2, 3). Wepresent paired d13Ccarb and d13Corg data from

1Department of Geosciences, Princeton University, Princeton,NJ 08544, USA. 2Geology and Geophysics, University of Ade-laide Mawson Laboratories, Adelaide, SA 5005, Australia.3Department of Earth and Planetary Sciences, NorthwesternUniversity, Evanston, IL 60208, USA.*Present address: Department of Earth and Planetary Sci-ences, McGill University, Montreal, Quebec H3A 2A7, Canada.†To whom correspondence should be addressed. E-mail:[email protected]

30 APRIL 2010 VOL 328 SCIENCE www.sciencemag.org608

REPORTS

on

Apr

il 29

, 201

0 w

ww

.sci

ence

mag

.org

Dow

nloa

ded

from

http://www.sciencemag.org

older Neoproterozoic strata: the Bitter SpringsFormation of the Amadeus Basin, central Aus-tralia, that was deposited during the Tonian Pe-riod (1000 to ~720 Ma) and the Etina-TrezonaFormations of the adjacent Adelaide Rift Com-plex deposited during the Cryogenian Period(~720 to 635 Ma) (5). Together with publisheddata from the Ediacaran Wonoka Formation (1),also of the Adelaide Rift Complex, these recordsspan numerous d13Ccarb shifts throughout theNeoproterozoic Era on a single continent, enabl-ing a test of the hypothesis that a large res-ervoir of DOC and corresponding invariancein d13Corg was a feature of the carbon cycle forthe entire era.

The paired carbon isotope data from the Bit-ter Springs Formation demonstrate covariationacross the onset of the ~800-Ma Bitter Springsstage and throughout the stage itself. At thetermination of the Bitter Springs stage, d13Ccarb

values shift abruptly from –2.7 per mil (‰) to+5.3‰, whereas d13Corg shifts from –29.9‰ to–26.7‰ (Fig. 1). The stable isotope results arevirtually identical between the 2-km-deep Wallaradrill core and a surface outcrop 120 km away,indicating that the signal is basin-wide and thatneither surface oxidation nor meteoric diagenesishave substantially altered the d13Ccarb–d

13Corg

record. The sympathetic shifts in d13Ccarb andd13Corg across the Bitter Springs stage confirmthat the stage reflects a large-scale perturbationto the isotopic composition of the DIC pool andthat organic matter in the sediments is repre-sentative of coeval biomass that fixed carbonfrom this 13C-depleted DIC (fig. S1). In starkcontrast, d13Corg values remain invariant acrossthe Cryogenian Trezona anomaly, in whichd13Ccarb drops by 18‰. Before the Trezonaanomaly, the d13Ccarb values of the Etina For-mation plateau at ~8‰, which is similar to thevalues observed in Cryogenian interglacial car-bonates from Namibia (6), Mongolia (7), andScotland (8). After deposition of the Enoramashale, carbonates of the subtidal Trezona For-mation record d13Ccarb values of –10‰ that in-crease up-stratigraphy to –2‰ before the glacialsediments of the Elatina Formation. Despite thesedramatic changes in d13Ccarb, d

13Corg values re-mained constant at –25‰ (Fig. 1). Unlikeduring the Shuram-Wonoka anomaly, therewas not an increase in the variability of d13Corg

values upwards through the Trezona anomaly.The new Australian data sets do not show sig-nificant correlation between d13Corg and totalorganic carbon, a proxy that is sometimes usedas evidence for alteration of the d13Corg signal(fig. S3). Taken together, the new data constrainthe buildup of a large DOC pool to after theBitter Springs stage of the mid-to-late Tonian butbefore the onset of the end-Cryogenian glaciation.

This timing for the growth of the DOC pooland the onset of non-steady-state dynamics isconsistent with very low sulfate levels in theCryogenian oceans (9) and a return to ferrugi-nous conditions in the deep ocean during early

Fig. 1. Carbonate carbon iso-tope values and organic carbonisotope values from Neoprotero-zoic carbonates of Australia withsimplified lithostratigraphy. Theplotted lithofacies represent thelithologies that dominate eachinterval, with wider boxes corre-sponding to deposition in shal-lower water. N389 field-sectionand Wallara-1 core data are fromthe Bitter Springs Formation ofthe Amadeus Basin, depositedbefore the Sturtian glacial event.C227 and C215 are part of a con-tinuous field section through theinterglacial stratigraphy of theAdelaide Rift Complex. Data forthe Wonoka Formation are from(1). The Bitter Springs, Trezona,and Shuram carbon isotope anom-alies are labeled next to thecarbon isotope records. The pre–Sturtian Islay anomaly is notrecorded in the Bitter SpringsFormation because of a discon-formity at the contact with theoverlying glacial sediments. TheGlobal Boundary Stratotype Sec-tion and Point for the Cryogenian/Ediacaran Period boundary is atthe contact between the glaci-genic Elatina formation and theoverlying Nuccaleena cap carbon-ate. Though it has yet to beformally defined with a strato-type section, we place the Tonian/Cryogenian boundary at the low-ermost evidence for Neoprotero-zoic glaciation, in concurrencewith the 2009 recommendationof the International Commissionon Stratigraphy.

Wal

lara

-1 c

ore

dept

h (m

)

AREY-ONGA

TREZ

ONA

Fm.

ETIN

A Fm

.Bit

ter Sp

rings

Fm.

ELAT

INAFm

.

C22

7 st

ratig

raph

ic h

eigh

t (m

)N

389

stra

tigra

phic

he

ight

(m

)C

215

stra

tigra

phic

he

ight

(m

)

LITHOFACIEScarbonates siliciclastics

stromatoliteribbonite

sandstonesiltstone

grainstone diamictite

S

M

δ13Ccarb

δ C

δ C

δ13Corg

δ13Ccarbδ13Corg

0

0

100

200

300

400

500

1300

1500

–1900

–1800

–1700

–1600

–1500

–2000

ED

IAC

AR

AN

CR

YO

GE

NIA

NT

ON

IAN

Bit

ter

Sp

rin

gs

Sta

ge

M SMarinoan glaciation Sturtian glaciation

unconformity

Trez

on

a A

no

mal

yS

hu

ram

An

om

aly

ENORAMA Fm.

BRACHINA, ABC, BUNYEROO Fms.

1400

600

700

800

–10 0 10–30 –20

–10 0 10–30 –20

–10 0 10–30 –20

200

100

WONO

KA F

m.0

100

200

300

400

500

600

700

Bun

yero

o G

orge

st

ratig

raph

ic h

eigh

t (m

)

~635 Ma

~551 Ma

~720 Ma

~800 Ma

HEAVI-TREE

this study Calver(2000)

POUN

DFm

.

Ediacaran Fossils

www.sciencemag.org SCIENCE VOL 328 30 APRIL 2010 609

REPORTS

on

Apr

il 29

, 201

0 w

ww

.sci

ence

mag

.org

Dow

nloa

ded

from


Cryogenian glaciation (10, 11). In such an ocean,decreased rates of aerobic respiration and bacte-rial sulfate reduction would slow organic carbonremineralization and extend the residence timeof DOC. Preferential remineralization of organicP over C, as occurs with increasing depth in themodern ocean (12), could increase the C:P ratioof DOC with the liberated phosphate, helping tosustain the productivity necessary to accumulatea large DOC pool. Furthermore, anoxic bottom-water conditions would favor burial of high C:Porganic matter because of decreased burial of Pbound to Fe oxyhydroxides (13), further main-taining nutrient supply and sustaining the pri-mary productivity required to explain elevatedCryogenian d13Ccarb values (14).

The return to ferruginous ocean conditionsand the buildup of DOC can be explained as aconsequence of global glaciation. The develop-ment of a thick continental regolith of uncon-solidated, chemically leached debris and soilduring the 1.5 billion years between the ~2.2-billion-year-old Paleoproterozoic Makganyeneglaciation of South Africa (15) and the ~720-Maearly Cryogenian pan-glacial event (16) wouldhave suppressed the Fe:S ratio of continentalrunoff. After reaching a depth of ~0.5 m, thethickness of regolith is inversely proportional tothe weatherability of the top of bedrock (Fig.2B) (17, 18). Although the average concentra-tions of Fe oxides are similar between sedimen-tary rocks and the rest of the upper continentalcrust, the average concentration of S is eight timesgreater in sedimentary lithologies (19). Thus, thedevelopment of a thick regolith on continentalinteriors in the absence of glacial erosion and

the preferential weathering of S-rich sedimenta-ry and ophiolitic rocks on continental margins,where tectonic uplift could facilitate physical re-moval of regolith, would have limited relative Feinput to the ocean. This mechanism for main-taining high relative S delivery helps to explainevidence for widespread euxinic conditionsthrough the Mesoproterozoic [1.6 to 1.0 billionyears ago (20, 21)].

The association of banded-iron formation(BIF) with Sturtian-age glacial deposits demon-strates that during the glaciation, Fe supply fromhydrothermal and continental-weathering sourcesexceeded sulfide availability (9). This Fe inputremoved available oxidants, resulting in anoxia,low sulfate levels, and BIF deposition. Althoughthe presence of BIF associated with the glacia-tion indicates transient ferruginous conditions,the maintenance of iron-rich deep oceans (10, 11)requires that a high relative flux of Fe continuedin the post-glacial period. Ubiquitous continentalice sheets during the Sturtian glaciation wouldhave scoured continental interiors, removingthe thick mantle of regolith. The relatively thickSturtian glacial deposits may represent physicalevidence of redeposited regolith that was erodedby dynamic early Cryogenian ice sheets, where-as the relatively thin Marinoan glacial depositsmay reflect the activity of stable late Cryogenianice sheets frozen to scoured bedrock—similar tothe Pleistocene evolution of the Canadian Shieldand Laurentide ice sheet (22). When ice sheetsretreated during the high CO2 escape from theearly Cryogenian glaciation, the vigorous weath-ering of freshly exposed continental crust wouldresult in a higher proportional delivery of Fe to

S into the ocean than during the preceding 1.5billion years. This postulated increase in therelative delivery of material derived from con-tinental interiors as compared with continentalmargins is supported by a steady increase in the87Sr/86Sr composition of the ocean after theSturtian ice age (23). Because the DOC reser-voir does not build up until after the Sturtianglaciation in this model, it predicts that d13Corg

will vary across the immediately pre-SturtianIslay negative d13Ccarb anomaly (8).

The sudden post-Sturtian increase in weath-erability could have led to lower equilibrium at-mospheric CO2 during the Cryogenian withoutchanges to volcanic CO2 input (Fig. 2C). Changesin CO2 are connected to the evolving sensitivityof silicate weathering rates to CO2 [weather-ability (kw)] and the varying fraction of totalcarbon burial that occurs as organic carbon ( forg)(14). High steady-state values of d13Ccarb duringthe late Tonian have been used to argue that ahigh forg helped lower CO2 before glaciation(14). The prevalence of continental landmass atlow latitude that facilitated high forg may haveincreased kw because of the abundance of silicaterocks associated with Grenville-age orogenicbelts in tropical weathering regimes. Landscapedisequilibrium associated with Bitter Springs–stage rapid true-polar wander (24) and increaseddelivery of moisture to continental interiors dur-ing the opening of incipient ocean basins asRodinia rifted apart (25) would have furtherincreased kw. Together, these late Tonian changeswould have reduced CO2 enough to initiateSturtian glaciation. However, it was the Sturtianglaciers themselves that scoured the continents,removing the long-lived Proterozoic regolith,greatly increasing continental weatherability,and setting up a new climatic regime with lowerCO2, a ferruginous ocean with high d13Ccarb,and a large DOC pool.

Although the close association of the Trezonaanomaly below Marinoan glacial deposits hasbeen interpreted as evidence for a causal rela-tionship between the two (14), the glacioeustaticsea level fall related to Marinoan glaciation didnot occur until after recovery from the most nega-tive d13Ccarb values. In some sections, TrezonaFormation d13Ccarb values recover to ~0‰ andare followed by more than 100 m of shallowing-upward peritidal sandstones before the first gla-cial deposits, further attenuating the connectionbetween the Trezona anomaly and glaciation.If the increase in kw and sustained high forg ofthe Cryogenian led to global cooling and oxy-gen release, the Trezona anomaly could reflectoxygenation of the deep ocean and partial re-mineralization of the large 13C-depleted DOCpool, as has been suggested for the Shuram-Wonoka anomaly. Organic carbon remineral-ization represents a negative climate feedback,releasing CO2 and preventing glaciation—whichis consistent with the lack of glacioeustatic changeduring the Trezona anomaly itself. The d13Ccarb

recovery and eventual Marinoan glaciation oc-

0 1 2 3

Weatherability of Continental

Interiors

Deep OceanChemistry

Carbon Cycle

LIP

regolith thickness (meters)

pCO2

loca

l wea

ther

abili

ty

NeoproterozoicMesoproterozoicPaleoproterozoic2.0 1.5 1.0

S-richFe-rich Fe

O2

high

low

little DOClarge DOC

Makg

anye

ne

Stu

rtia

n

Marin

oan

Gas

kier

s

high

kw

, hig

h f or

g

(Cry

ogen

ian)

low kw, low forg

(Mesoproterozoic)

med

ium k w, h

igh for

g

(Ton

ian)

A B

C

in=

F out

)

Fig. 2. (A) Summary illustration relating Proterozoic deep-ocean chemistry [modified from (26)], smallversus large DOC carbon cycle, and the weatherability of continental interiors. Before the first Neo-proterozoic glaciation, there was no unusual concentration of large igneous province events [shown ascompiled by (27), with solid and dashed lines indicating <20- and >20-Ma uncertainty, respectively].Because local weatherability is a function of regolith thickness [(B), modified from (18)], regolithdevelopment on continental interiors through the nonglacial Mesoproterozoic (28) would lead to thedepicted decrease in regional weatherability in (A). The relationship shown schematically in (C) is Fin =(kw × MCO2 )/(1 – forg), where kw is the slope of the weathering-CO2 feedback and is partly a function ofthe regional weatherability depicted in (B). An increase in kw and in the relative burial of C as organicmatter can result in a decrease in CO2, as shown for the Mesoproterozoic → Tonian → Cryogenian,without changes in volcanic An increase input.

30 APRIL 2010 VOL 328 SCIENCE www.sciencemag.org610

REPORTS

on

Apr

il 29

, 201

0 w

ww

.sci

ence

mag

.org

Dow

nloa

ded

from


curred when the large DOC pool had been re-duced in size enough to no longer represent anegative feedback to global climatic cooling.

References and Notes1. C. R. Calver, Precambrian Res. 100, 121 (2000).2. D. A. Fike, J. P. Grotzinger, L. M. Pratt, R. E. Summons,

Nature 444, 744 (2006).3. K. A. McFadden et al., Proc. Natl. Acad. Sci. U.S.A. 105,

3197 (2008).4. D. H. Rothman, J. M. Hayes, R. E. Summons, Proc. Natl.

Acad. Sci. U.S.A. 100, 8124 (2003).5. Data tables and methods are available as supporting

material on Science Online.6. G. P. Halverson, P. F. Hoffman, D. P. Schrag, A. C. Maloof,

A. H. N. Rice, Geol. Soc. Am. Bull. 117, 1181 (2005).7. F. A. Macdonald, D. S. Jones, D. P. Schrag, Geology 37,

123 (2009).8. A. R. Prave, A. E. Fallick, C. W. Thomas, C. M. Graham,

J. Geol. Soc. London 166, 845 (2009).9. M. T. Hurtgen, M. A. Arthur, N. Suits, A. J. Kaufman,

Earth Planet. Sci. Lett. 203, 413 (2002).10. D. E. Canfield et al., Science 321, 949 (2008).

11. C. Li et al., Science 328, 80 (2010).12. P. Sannigrahi, E. D. Ingall, R. Benner, Geochim.

Cosmochim. Acta 70, 5868 (2006).13. P. Van Cappellen, E. D. Ingall, Paleoceanography 9, 677

(1994).14. D. P. Schrag, R. A. Berner, P. F. Hoffman, G. P. Halverson,

Geochem. Geophys. Geosyst. 10, 1029/2001GC000219(2002).

15. R. E. Kopp, J. L. Kirschvink, I. A. Hilburn, C. Z. Nash,Proc. Natl. Acad. Sci. U.S.A. 102, 11131 (2005).

16. F. A. Macdonald et al., Science 327, 1241 (2010).17. F. Ahnert, Ed., Geomorphological Models: Theoretical

and Empirical Aspects (Catena, Reiskirchen, Germany,1987), pp. 31–50.

18. E. J. Gabet, S. M. Mudd, Geology 37, 151 (2009).19. K. H. Wedepohl, Geochim. Cosmochim. Acta 59, 1217

(1995).20. D. E. Canfield, Nature 396, 450 (1998).21. T. W. Lyons, A. D. Anbar, S. Severmann, C. Scott,

B. C. Gill, Annu. Rev. Earth Planet. Sci. 37, 507 (2009).22. P. U. Clark, D. Pollard, Paleoceanography 13, 1 (1998).23. G. P. Halverson, F. O. Dudas, A. C. Maloof, S. A. Bowring,

Palaeogeogr. Palaeoclimatol. Palaeoecol. 256, 103 (2007).24. A. C. Maloof et al., Geol. Soc. Am. Bull. 118, 1099 (2006).

25. Y. Goddéris et al., C. R. Geosci. 339, 212 (2007).26. D. T. Johnston, F. Wolfe-Simon, A. Pearson, A. H. Knoll,

Proc. Natl. Acad. Sci. U.S.A. 106, 16925 (2009).27. R. E. Ernst, K. L. Buchan, Spec. Pap. Geol. Soc. Am. 352,

483 (2001).28. L. C. Kah, R. Riding, Geology 35, 799 (2007).29. We thank K. Bovee, R. Levin, W. Jacobsen, L. Wingate,

and L. Godfrey for assistance with sample preparationand analysis and D. Rothman, L. Kah, R. Kopp, N. Cassar,J. Higgins, J. Husson, and D. Sigman for comments anddiscussions. This work was supported by NSF grantsEAR-0514657 and EAR-084294 to A.C.M., EAR-0720045to M.T.H., an American Association of PetroleumGeologists Grant to C.V.R., and an NSF East Asia andPacific Summer Institute fellowship to N.L.S.-H.

Supporting Online Materialwww.sciencemag.org/cgi/content/full/328/5978/608/DC1Materials and MethodsFigs. S1 to S4Table S1References

10 November 2009; accepted 11 March 201010.1126/science.1184508

Asian Monsoon Transport of Pollutionto the StratosphereWilliam J. Randel,1* Mijeong Park,1 Louisa Emmons,1 Doug Kinnison,1 Peter Bernath,2,3Kaley A. Walker,4,3 Chris Boone,3 Hugh Pumphrey5

Transport of air from the troposphere to the stratosphere occurs primarily in the tropics, associated withthe ascending branch of the Brewer-Dobson circulation. Here, we identify the transport of air massesfrom the surface, through the Asian monsoon, and deep into the stratosphere, using satellite observationsof hydrogen cyanide (HCN), a tropospheric pollutant produced in biomass burning. A key factor inthis identification is that HCN has a strong sink from contact with the ocean; much of the air in the tropicalupper troposphere is relatively depleted in HCN, and hence, broad tropical upwelling cannot be themain source for the stratosphere. The monsoon circulation provides an effective pathway for pollutionfrom Asia, India, and Indonesia to enter the global stratosphere.

The Asian summer monsoon circulationcontains a strong anticyclonic vortex inthe upper troposphere and lower strato-

sphere (UTLS), spanning Asia to the MiddleEast. The anticyclone is a region of persistent

enhanced pollution in the upper troposphereduring boreal summer, linked to rapid verticaltransport of surface air from Asia, India, andIndonesia in deep convection, and confinementby the strong anticyclonic circulation (1–6). A

mean upward circulation on the eastern side ofthe anticyclone extends the transport into thelower stratosphere, as evidenced by satelliteobservations of water vapor (7) and ozone (8),plus carbon monoxide and other pollution tracers(1, 4, 5). Model calculations have suggested thattransport from the monsoon region could con-tribute substantially to the budget of stratosphericwater vapor (8, 9), but this effect has not beenisolated from broader-scale tropical upwelling inobservational data.

Hydrogen cyanide (HCN) is produced pri-marily as a result of biomass and biofuel burningand is often used as a tracer of pollution originat-ing from fires (10–12). In the free atmosphere,

1National Center for Atmospheric Research, Boulder, CO, USA.2Department of Chemistry, University of York, Heslington, UK.3Department of Chemistry, University of Waterloo, Waterloo,Ontario, Canada. 4Department of Physics, University ofToronto, Toronto, Ontario, Canada. 5School of GeoSciences,University of Edinburgh, Edinburgh, UK.

*To whom correspondence should be addressed. E-mail:[email protected]

Fig. 1. Time average mixing ratio [parts per billion by volume (ppbv)] ofHCN near 13.5 km during boreal summer (June to August) derived from (A)ACE-FTS observations and (B) WACCM chemical transport model calcu-

lations. Arrows in both panels denote winds at this level derived frommeteorological analysis, showing that the HCN maximum is linked with theupper tropospheric Asian monsoon anticyclone.

www.sciencemag.org SCIENCE VOL 328 30 APRIL 2010 611

REPORTS

on

Apr

il 29

, 201

0 w

ww

.sci

ence

mag

.org

Dow

nloa

ded

from


www.sciencemag.org/cgi/content/full/328/5978/608/DC1

Supporting Online Material for

Cryogenian Glaciation and the Onset of Carbon-Isotope Decoupling

Nicholas L. Swanson-Hysell, Catherine V. Rose, Claire C. Calmet, Galen P. Halverson, Matthew T. Hurtgen, Adam C. Maloof*

*To whom correspondence should be addressed. E-mail: [email protected]

Published 30 April 2010, Science 328, 608 (2010)

DOI: 10.1126/science.1184508

This PDF file includes:

Materials and Methods

Figs. S1 to S4

Table S1

References

Supplementary Online Materials: Cryogenian glaciation and the onset of carbon-isotope decoupling

S1 Methods

S1.1 δ13Ccarb methods

All samples used in this study were collected in the course of measuring stratigraphic sections or logging

drill core. Samples were chosen so as to minimize veining, fractures, and siliciclastic material. In order to

prepare powders for carbonate carbon isotopic analysis, samples were slabbed perpendicular to laminations,

polished to clarify internal structure and subsampled with 1 mm dental drill bits. At the University of

Michigan Stable Isotope Laboratory, all powders were heated to 200◦C to remove volatile contaminants and

water. Samples were then placed in individual borosilicate reaction vessels and reacted at 76◦C with 3 drops

of H3PO4 in a Finnigan MAT Kiel I preparation device coupled directly to the inlet of a Finnigan MAT 251

triple collector isotope ratio mass spectrometer. δ13C and δ18O data were acquired simultaneously and are

reported in the standard delta notation as the h difference from the VPDB standard. Precision and accuracy

of data are monitored through daily analysis of at least six standards which are run to bracket sample suites

at the beginning, middle, and end of the day’s run. Measured precision is maintained at better than 0.1h

(1σ) for both δ13C and δ18O.

S1.2 δ13Corg methods

Organic carbon isotopic values were obtained from the total organic carbon (TOC) of insoluble residues.

After removing the outside layer of surface oxidation and large veins, whole rock samples were crushed

into powder. Insoluble residues for organic carbon isotope analysis were obtained by acidifying these whole

rock powders in 6N HCl for 24 hours to dissolve all carbonate minerals. Care was taken to ensure that

acid was added and acidification continued until there was absolutely no visible carbonate dissolution so

that the analyses would not be affected by contamination from residual inorganic carbon. The insoluble

residues were then rinsed with DI water, dried and loaded into tin capsules for isotopic analysis. At the

University of Adelaide, samples were flash combusted at 1030◦C in a Fisons Elemental Analyzer. The

1

resulting CO2 gas was analyzed by continuous flow on a Fisons Optima isotope ratio mass spectrometer.

δ13Corg values were calibrated against in-house glycine and glutamic acid standards with bracketing isotopic

values. Reproducibility (typically better than 0.5h) was verified by duplicate sample analyses and regularly

interspersed in-house sucrose standards (δ13Corg = -25.8h). Values are reported in standard delta notation

relative to VPDB. At Rutgers University, δ13Corg values were obtained on a GVI Isoprime CF-IRMS linked

to a Eurovector elemental analyzer. Isotope ratios were corrected against NBS 22 using the accepted value of

-30.03h (S1). Organic C concentrations were measured using standards with known carbon concentration

and the intensity of masses 44 and 28. Isotope and concentration standards were run following eight sample

unknowns. TOC values for the bulk samples were calculated by combining the carbon concentration data

with measurements of the ratio of insoluble residue to original pre-decarbonated powder.

S2 Data

The isotopic data and total organic carbon values for the N389 field section (23◦31’7"S 134◦26’55.03"E),

C215 and C227 field sections (31◦23’44”S, 138◦51’14”E), and the Wallara-1 stratigraphic drill core (24◦36’55”S,

132◦20’23”E) are presented in Table S1. Previously developed carbon isotope data for the Wallara-1 strati-

graphic drill core show similar trends and absolute values to the new data presented here, but at much lower

resolution (S2). Cross plots of the data from Figure 1 are shown in Figures S1-S4. These plots illustrate

some important features of the data sets.

• Sympathetic shifts between δ13Ccarb and δ13Corg for the Bitter Springs Formation data can be seen

visually in Fig. S1 and result in a high R2 value (0.71) in contrast with the low R2 value (0.22) for the

combine Etina/Trezona results (Fig. S1).

• In the ∆δ13C vs. δ13Ccarb cross plots presented in Fig. S2 the fits to the data from the Trezona Forma-

tion (along with the Wonoka, Shuram and Doushantuo) have high R2 values and slopes approaching

unity. This high correlation combined with the slopes of ∼1 can be explained as a result of the vari-

ability in ∆δ13C primarily being attributable to variation in δ13Ccarb.

• The plots of δ13Corg vs. total organic carbon (TOC) % show that there is no observed dependence

2

between the values obtained for the isotopes of the organic matter in the samples and the concentration

of that organic matter.

• Knauth and Kennedy (S3) proposed covariation between δ13Ccarb and δ18Ocarb as an index for differ-

entiating between original and altered δ13Ccarb values. The δ18Ocarb values of the Etina and Trezona

Formations are scattered around -10h despite the very large difference in δ13Ccarb values. In the

Bitter Springs data there is no covariation between the large swings in δ13Ccarb values and the scatter

in δ18Ocarb values.

Supplemental References

[S1] T. Coplen, et al., Analytical Chemistry 78, 2439 (2006).

[S2] A. C. Hill, K. Arouri, P. Gorjan, M. R. Walter, in Carbonate Sedimentation and Diagenesis in an

Evolving Precambrian World (SEPM Special Publications, Tulsa, 2000), vol. 67, pp. 327–344.

[S3] L. P. Knauth, M. J. Kennedy, Nature 460, 728 (2009).

3

-35 -30 -25 -20-6

-4

-2

0

2

4

6

8

δ13Corg

δ13C ca

rb

-35 -30 -25-10

-8

-6

-4

-2

0

δ13Corg

δ13C ca

rb

-30 -25 -200

2

4

6

8

10

δ13Corg

δ13C ca

rb

-28 -26 -24 -22 -20 -18

4

6

8

10

12

δ13Corg

δ13C ca

rb

-30 -25 -20-12

-10

-8

-6

-4

-2

0

2

δ13Corg

δ13C ca

rb

-35 -30 -25 -20 -15-10

-5

0

5

10

δ13Corg

δ13C ca

rb

-30 -25 -20 -15-10

-5

0

5

δ13Corg

δ13C ca

rb

-40 -35 -30 -25 -20 -15-15

-10

-5

0

5

10

δ13Corg

δ13C ca

rb

-35 -30 -25 -20 -15 -10-10

-5

0

5

δ13Corg

δ13C ca

rb

syn-Bitter Springs Stage(Wallara1)

post-Bitter Springs Stage(Wallara1)

all Bitter Spring Formation(Wallara1)

Trezona Formation (C215)Etina/Trezona Formation(C215, C227 combined)Etina Formation (C227)

Shuram Formation Doushantuo FormationWonoka Formation

R2=0.13 R2=0.02 R2=0.22

R2=0.17 R2=0.27 R2=0.71

R2=0.06 R2=0.42 R2=0.01

Figure S1: δ13Corg vs. δ13Ccarb for the data presented here from the Bitter Springs, Etina and Trezona Formations as well as

data from the Wonoka Formation (1), the Shuram Formation (2) and the Doushantuo Formation (3) with calculated R2 values.

The syn-Bitter Springs Stage plot is of data from between meter levels 1920 to 1795 of the Wallara-1 stratigraphic drill core while

post-Bitter Springs Stage data is from between meter levels 1790 to 1425. Sample C215-148.3 was excluded as its δ13Corg value

was anomalous in comparison to nearby samples.

4

20 25 30 35 40-10

-5

0

5

10

Δδ13C

δ13C ca

rb

24 26 28 30

-6

-4

-2

0

Δδ13C

δ13C ca

rb

28 30 32 34

2

4

6

8

Δδ13C

δ13C ca

rb

25 30 35 400

5

10

15

Δδ13C

δ13C ca

rb

10 15 20 25-15

-10

-5

0

Δδ13C

δ13C ca

rb

10 20 30 40-15

-10

-5

0

5

10

15

Δδ13C

δ13C ca

rb

10 20 30 40-15

-10

-5

0

5

10

15

Δδ13C

δ13C ca

rb

10 20 30 40-15

-10

-5

0

5

10

15

Δδ13C

δ13C ca

rb

slope= 0.37 [0.29 0.57]intercept=-8.5 [-14.2 -6.2]R2=0.63

slope= 0.50 [-0.68 0.66]intercept=-16.6 [-21.0 15.7]R2=0.63



slope= 0.49 [-0.53 0.59]intercept=-7.5 [-10.9 25.4]R2=0.07

slope=0.84 [0.46 0.97]intercept=-20.4 [-14.7 -22.3]R2=0.87



pre-Bitter Springs Stage(N389)

Trezona Formation (C215)Etina Formation (C227)

Shuram Formation Doushantuo FormationWonoka Formation

syn-Bitter Springs Stage(Wallara1)

post-Bitter Springs Stage(Wallara1)

Figure S2: ∆δ13C vs. δ13Ccarb for the data presented here from the Bitter Springs, Etina and Trezona Formations as well as data

from the Wonoka Formation (1), the Shuram Formation (2) and the Doushantuo Formation (3). The linear fits were computed with

the reduced major axis method and the 95% confidence intervals presented were obtained from 1000 bootstrapped data sets. R2

values are also presented for each data set. The syn-Bitter Springs Stage plot is of data from between meter levels 1920 to 1795

of the Wallara-1 stratigraphic drill core while post-Bitter Springs Stage data is from between meter levels 1790 to 1425. Samples

C227-146.6 and C227-718.0 were excluded from the fit calculated for the Etina data due to anomalous δ13Ccarb values.

5

-35 -30 -25 -20 -150

0.005

0.01

0.015

0.02

0.025

0.03

0.035

0.04

δ13Corg

TOC

(%)

-35 -30 -25 -20 -150

0.005

0.01

0.015

0.02

0.025

0.03

0.035

0.04

δ13Corg

TOC

(%)

-35 -30 -25 -20 -150

0.02

0.04

0.06

0.08

0.1

δ13Corg

TOC

(%)

-35 -30 -25 -20 -150

0.02

0.04

0.06

0.08

0.1

δ13Corg

TOC

(%)

R2=0.009 R2=0.001

R2=0.048 R2=0.305


Bitter Springs Formation (N389) Bitter Springs Formation (Wallara1)

Figure S3: Plots of δ13Corg vs. total organic carbon (TOC) %. There is no observed dependence between TOC and the isotopic

values of the organic matter.

6

-15 -10 -5 0 5-10

-5

0

5

10

δ18Ocarb

δ13C ca

rb

-15 -10 -5 0 5-10

-5

0

5

10

δ18Ocarb

δ13C ca

rb

-15 -10 -5 0 5-10

-5

0

5

10

δ18Ocarb

δ13C ca

rb

-15 -10 -5 0 5-10

-5

0

5

10

δ18Ocarb

δ13C ca

rb


Bitter Springs Formation (N389) Bitter Springs Formation (Wallara1)

Figure S4: Plots of δ18Ocarb vs. δ13Ccarb. The variability between the high δ13Ccarb values of the Etina Formation and the low

δ13Ccarb values of the Trezona Formation is not accompanied by a notable shift in δ18Ocarb. Within the Bitter Springs formation

the large shifts in δ13Ccarb show no covariance with δ18Ocarb.

7

Table S1: Isotopic data from the Wallara-1 drill core, P8 field section, C215 field section,

C227 field section

section/core stratigraphic level δ13Ccarb δ18Ocarb δ13Corg ∆δ13C TOC % Corg lab sample type

Wallara-1 2000.9∗∗ -1.0 -6.1 dolostone

Wallara-1 2000.8 -1.1 -4.9 dolostone

Wallara-1 2000.7 -27.6 26.5 Adelaide insoluble residue




Wallara-1 1996.7 -1.0 -5.3 -29.0 27.9 0.036 Rutgers dolostone






Wallara-1 1991.5 0.1 -5.3 dolostone



Wallara-1 1988.6 -0.8 -5.2 -30.0 29.3 Adelaide dolostone






Wallara-1 1981.9†† 0.6 -4.5 -27.7 28.3 0.016 Rutgers dolostone










Wallara-1 1973.8∗∗ -28.0 28.8 Adelaide insoluble residue





8


C227 field section cont.









Wallara-1 1966.0∗∗ -1.3 -5.6 dolostone







Wallara-1 1960.0 1.9 -2.2 -27.6 29.5 Adelaide dolostone





Wallara-1 1955.3 -28.3 32.4 0.109 Rutgers insoluble residue


Wallara-1 1953.5 4.2 -3.4 -27.0 0.014 Rutgers dolostone

Wallara-1 1952.0 4.8 -5.0 -26.6 31.4 dolostone



Wallara-1 1950.1∗∗ 4.4 -4.4 dolostone




Wallara-1 1946.1∗∗ 6.4 -5.8 -27.3 33.7 0.014 Rutgers dolostone


Wallara-1 1937.0 1.3 -5.2 -29.1 0.076 Rutgers dolostone

Wallara-1 1936.3 -27.5 28.8 insoluble residue





9









































10
























Wallara-1 1882.0∗∗∗ -2.9 -9.8 limestone







Wallara-1 1876.4∗∗ -3.8 -5.3 -30.9 27.1 Adelaide dolostone










11




































Wallara-1 1839.6 -3.1 -4.1 -31.0 27.9 Adelaide dolostone





12






















Wallara-1 1820.9∗∗ -2.5 -4.8 -30.4 27.9 Adelaide dolostone



















13









Wallara-1 1793.0 -1.5 -4.3 -26.8 25.3 0.007 Adelaide dolostone









Wallara-1 1741.3 5.0 -5.9 limestone























14


























Wallara-1 1671.1†† 3.7 -7.7 -25.3 29.0 0.006 Rutgers limestone



Wallara-1 1665.2 4.3 -8.2 -25.8 30.1 0.006 Adelaide limestone




Wallara-1 1659.8∗∗∗ -25.8 31.9 Adelaide insoluble residue







Wallara-1 1648.4 5.8 -4.8 -25.1 30.9 0.021 Rutgers dolostone

15









Wallara-1 1640.5∗∗∗ 6.8 -6.9 limestone







Wallara-1 1634.9∗∗∗∗ -24.6 31.3 Adelaide insoluble residue






















Wallara-1 1615.9 5.5 -7.9 -27.2 32.7 0.041 Rutgers limestone



16









































17









































18








Wallara-1 1434.1 -26.4 30.6 insoluble residue













Wallara-1 1425.4∗∗ 3.2 -4.5 dolostone

Wallara-1 1424.3∗∗ 4.3 0.5 dolostone

N389 33.4 1.9 -3.5 -26.6 28.6 0.034 Adelaide dolostone

N389 34.6 2.0 -3.8 dolostone

N389 59.2 2.6 -2.8 dolostone

N389 59.6 1.1 -2.6 dolostone

N389 61.5†† 2.1 -2.8 -29.2 31.2 0.028 Adelaide dolostone

N389 62.5∗∗ 1.9 -2.7 dolostone

N389 64.1 2.6 -2.3 dolostone

N389 65.3 2.2 -2.7 dolostone

N389 66.1∗∗ 1.7 -2.3 dolostone

N389 68.2 2.2 -3.0 dolostone

N389 69.7 3.1 -3.0 dolostone

N389 70.8 2.9 -2.6 dolostone

N389 72.7 2.1 -3.5 dolostone

N389 75.0 1.8 -2.7 dolostone

N389 76.4∗∗ 1.9 -2.7 dolostone

N389 77.1 1.6 -2.4 dolostone


N389 78.3 1.1 -3.0 dolostone

19




N389 79.8∗∗ 0.0 -3.7 dolostone

N389 83.4 0.4 -2.3 dolostone

N389 85.4 1.3 -2.3 dolostone


N389 87.3 -0.0 -2.9 dolostone

N389 87.9 0.0 -2.7 dolostone

N389 89.3 0.4 -3.0 dolostone

N389 88.4 3.2 -2.8 dolostone

N389 90.4 0.9 -1.9 dolostone

N389 91.0 0.6 -2.6 dolostone

N389 92.3 0.8 -2.8 dolostone

N389 94.4∗∗ 1.1 -3.0 dolostone


N389 98.2 1.4 -2.4 dolostone

N389 99.6 2.2 -2.1 dolostone

N389 103.4 2.3 -2.1 dolostone

N389 106.2 1.1 -5.0 dolostone

N389 109.0 1.9 -2.7 dolostone

N389 111.0 1.7 -3.4 dolostone

N389 114.3 2.9 -2.8 dolostone


N389 117.6∗∗ 2.1 -2.5 dolostone

N389 119.2∗∗ 2.7 -2.5 dolostone

N389 120.8 2.3 -2.6 dolostone

N389 123.2 1.5 -2.3 dolostone

N389 124.2 0.6 -2.0 dolostone

N389 139.0 1.9 -2.8 dolostone

N389 141.1 3.0 -2.1 dolostone


N389 143.7 2.7 -2.5 dolostone

N389 145.8 3.5 -3.3 dolostone

N389 148.3 2.9 -1.9 dolostone

N389 149.9∗∗ 3.2 -2.1 dolostone

N389 152.3 2.8 -2.2 dolostone

N389 153.4 2.6 -2.2 dolostone

N389 155.4 3.2 -2.5 dolostone


20




N389 157.8 3.0 -2.1 dolostone

N389 158.9 3.0 -2.4 dolostone

N389 161.2∗∗ 3.0 -1.9 dolostone

N389 162.7 2.8 -2.3 dolostone

N389 164.1 3.4 -1.8 dolostone

N389 165.9 4.0 -1.7 dolostone

N389 168.8 3.2 -1.3 dolostone

N389 170.1 2.6 -2.2 dolostone


N389 173.7 1.8 -1.8 dolostone

N389 176.9 2.1 -1.6 dolostone

N389 179.0 2.8 -1.8 dolostone

N389 180.4 1.7 -1.5 dolostone


N389 184.3 3.0 -1.5 dolostone

N389 186.3∗∗ 2.8 -2.5 dolostone

N389 187.7 2.9 -1.1 dolostone

N389 188.4 3.1 -2.7 dolostone


N389 193.1 3.5 -1.5 dolostone

N389 195.4 3.4 -1.4 dolostone

N389 198.2 2.8 -1.6 dolostone

N389 200.3 3.0 -1.7 dolostone

N389 201.3 2.8 -1.5 dolostone

N389 203.4 2.2 -1.8 dolostone

N389 204.5 2.1 -1.5 dolostone

N389 207.0 1.9 -1.3 dolostone

N389 208.9 1.2 -1.4 dolostone


N389 211.6 -0.2 -1.9 dolostone

N389 214.7 0.7 -1.8 dolostone

N389 215.7 -0.4 -3.6 dolostone


N389 219.0 1.2 -1.2 dolostone

N389 219.0 1.2 -1.2 dolostone

N389 221.2 -0.0 -3.0 dolostone

N389 222.8 2.1 -1.3 dolostone

21




N389 224.4 2.5 -0.7 dolostone

N389 225.4 2.5 -1.1 dolostone

N389 226.5∗∗ 2.6 -1.1 dolostone

N389 228.4∗∗†† 2.8 -1.3 -28.2 31.0 0.040 Adelaide dolostone

N389 230.3 2.5 -1.8 dolostone

N389 232.3 2.9 -1.0 dolostone

N389 235.0 3.0 -0.7 dolostone


N389 238.2 2.7 -1.5 dolostone

N389 239.9 3.2 -1.6 dolostone

N389 241.8 3.2 -1.6 dolostone


N389 244.6 2.6 -2.1 dolostone

N389 245.6 3.1 -2.1 dolostone

N389 248.2 3.1 -2.1 dolostone

N389 253.2 3.9 -2.2 dolostone

N389 257.6 2.6 -2.4 dolostone

N389 258.3 2.0 -2.7 dolostone

N389 264.9 4.2 -1.8 dolostone

N389 267.0 4.4 -2.2 dolostone

N389 267.7∗∗ 4.5 -1.9 -32.3 36.8 0.059 Adelaide dolostone

C215 0.3∗∗ -8.7 -13.0 limestone

C215 0.5 -8.6 -12.6 limestone

C215 12.3 -8.8 -12.5 -23.9 15.1 0.011 Rutgers limestone

C215 13.9 -8.7 -11.7 limestone

C215 14.9 -8.6 -12.3 limestone

C215 14.9 -8.5 -12.8 limestone

C215 15.5 -8.5 -12.5 limestone

C215 16.2 -8.4 -12.7 limestone

C215 17.1∗∗ -8.6 -12.9 limestone


C215 27.4 -8.6 -12.4 limestone

C215 34.1 -8.2 -12.6 limestone

C215 34.5 -8.2 -12.6 limestone


C215 36.1 -8.6 -11.8 limestone

C215 36.4 -8.7 -12.6 limestone

22




C215 38.3∗∗ -9.5 -12.1 -24.4 14.9 0.016 Rutgers limestone

C215 46.5∗∗ -8.5 -12.8 limestone

C215 47.7 -8.7 -12.1 limestone

C215 48.3 -8.8 -12.6 limestone

C215 54.8 -8.4 -11.8 limestone

C215 55.9 -8.6 -12.2 limestone

C215 56.2 -8.8 -11.2 limestone

C215 56.2 -8.3 -12.2 limestone

C215 56.3 -8.8 -12.5 limestone

C215 57.7 -8.3 -12.4 limestone

C215 59.4∗∗ -9.0 -11.9 limestone

C215 61.7 -8.6 -12.1 limestone

C215 63.4∗∗ -9.7 -12.3 limestone

C215 63.8 -8.8 -12.2 limestone


C215 67.1 -8.6 -12.2 limestone

C215 67.9 -8.5 -12.1 limestone

C215 68.6∗∗ -8.8 -11.6 limestone


C215 72.9 -8.7 -12.1 limestone

C215 73.3 -8.3 -12.0 limestone

C215 74.5∗∗ -8.2 -12.2 limestone

C215 75.3 -8.1 -11.8 limestone

C215 76.5 -8.2 -12.2 limestone


C215 79.5 -26.2 0.012 Rutgers limestone

C215 77.3 -8.5 -12.1 limestone

C215 80.5 -8.3 -11.2 limestone

C215 81.2∗∗ -8.9 -10.3 limestone

C215 81.5 -8.8 -11.9 limestone


C215 82.6 -9.4 -11.7 limestone

C215 87.3 -8.3 -11.7 limestone

C215 89.4 -8.9 -11.5 limestone

C215 91.5∗∗ -8.6 -11.3 limestone

C215 93.7 -8.7 -11.9 limestone

C215 94.6 -8.7 -11.9 limestone

23




C215 97.0∗∗ -8.6 -11.9 limestone

C215 98.2 -8.3 -11.7 limestone

C215 99.2 -8.1 -11.7 limestone

C215 100.1∗∗ -8.4 -11.8 limestone

C215 100.7 -9.2 -11.4 limestone


C215 103.8 -9.6 -11.6 limestone

C215 106.5∗∗ -8.1 -11.1 limestone

C215 107.4 -7.9 -10.6 limestone

C215 108.3 -8.5 -9.4 limestone


C215 117.8 -8.3 -11.4 limestone

C215 118.4 -8.1 -10.8 limestone

C215 122.4∗∗∗ -7.7 -11.3 limestone

C215 124.0∗∗ -7.5 -11.3 limestone

C215 124.8∗∗ -7.7 -11.2 limestone

C215 125.8∗∗ -7.5 -10.9 limestone

C215 126.6∗∗ -7.5 -9.5 limestone

C215 127.1††∗∗∗∗ -7.5 -8.9 -27.9 20.4 0.009 Rutgers limestone

C215 127.8∗∗ -7.7 -10.5 limestone

C215 128.6∗∗ -7.4 -8.5 limestone

C215 129.8∗∗ -7.6 -8.5 limestone

C215 130.8∗∗ -7.9 -9.6 limestone

C215 131.6∗∗ -8.0 -8.9 limestone

C215 132.5∗∗ -7.7 -8.7 limestone

C215 133.1∗∗∗∗ -7.7 -9.1 limestone

C215 134.2∗∗∗∗ -7.7 -8.6 limestone

C215 135.2∗∗ -7.4 -10.2 limestone

C215 136.0 -7.6 -10.4 limestone

C215 136.8 -7.6 -10.2 limestone


C215 138.1 -7.5 -9.1 limestone

C215 138.6 -7.2 -10.2 limestone


C215 142.7 -7.4 -10.0 limestone


C215 149.5 -7.2 -9.9 limestone

24




C215 150.2∗∗ -7.3 -9.9 limestone

C215 152.2 -7.1 -9.2 limestone

C215 153.8 -7.4 -8.8 limestone


C215 155.7 -7.4 -7.6 limestone

C215 156.4 -7.3 -9.2 limestone

C215 156.7∗∗ -6.8 -9.4 limestone

C215 157.2 -6.8 -9.3 limestone

C215 157.7 -6.7 -9.5 limestone

C215 158.5 -6.7 -9.6 limestone


C215 160.5 -6.6 -9.5 limestone

C215 161.0 -6.7 -9.3 limestone

C215 162.3 -6.4 -9.6 limestone

C215 162.9∗∗ -6.2 -9.6 limestone


C215 166.4 -6.5 -9.7 limestone

C215 167.3 -6.4 -9.5 limestone

C215 167.6 -6.3 -9.4 limestone


C215 169.8 -6.4 -9.5 limestone

C215 170.5∗∗ -6.4 -9.4 limestone

C215 171.4 -6.4 -9.6 limestone

C215 172.4 -6.1 -9.4 limestone

C215 173.1 -6.1 -9.6 limestone

C215 174.1 -6.1 -9.5 limestone

C215 175.1 -6.3 -9.7 limestone

C215 175.3 -6.0 -9.6 limestone

C215 175.4††††∗∗ -6.2 -9.5 -26.1 19.8 0.009 Rutgers limestone

C215 176.5 -6.0 -9.6 limestone

C215 177.8 -6.0 -9.6 limestone

C215 179.0∗∗ -5.8 -9.4 limestone

C215 180.6 -5.9 -9.6 limestone

C215 181.0 -5.9 -9.6 limestone

C215 181.3 -5.7 -8.7 limestone

C215 182.4 -5.9 -8.8 limestone

C215 183.6 -5.8 -8.6 limestone

25





C215 185.3 -6.2 -9.6 limestone

C215 186.3 -6.1 -9.4 limestone

C215 187.0 -6.1 -9.5 limestone


C215 189.7 -5.7 -9.6 limestone

C215 190.6 -5.5 -9.3 limestone

C215 191.7 -5.6 -9.5 limestone

C215 192.5 -5.8 -9.3 limestone

C215 193.0∗∗ -5.9 -9.4 limestone


C215 194.1 -5.7 -9.6 limestone

C215 195.1∗∗ -5.7 -8.1 limestone

C215 196.8†† -5.9 -9.3 -25.2 19.3 0.004 Rutgers limestone

C215 197.1 -5.5 -9.4 limestone

C215 197.6 -5.5 -9.6 limestone

C215 198.8 -5.5 -9.6 limestone

C215 199.7∗∗ -5.4 -8.9 limestone


C215 201.7 -5.3 -9.4 limestone

C215 201.9 -5.5 -9.1 limestone

C215 203.1 -5.2 -9.1 limestone

C215 204.1∗∗ -5.7 -8.7 limestone

C215 205.1 -5.6 -8.9 limestone

C215 206.6 -6.2 -8.2 limestone



C215 208.5 -5.4 -8.9 limestone

C215 209.6 -5.5 -9.3 limestone

C215 211.1 -4.7 -9.5 limestone

C215 212.1 -5.0 -10.0 limestone

C215 212.8 -4.8 -9.6 limestone

C215 213.5 -4.8 -9.9 limestone


C215 215.1 -4.8 -8.8 limestone

C215 216.1∗∗ -5.3 -9.8 limestone

C215 216.9 -5.3 -10.0 limestone

26





C215 219.8 -4.9 -9.8 limestone

C215 221.1 -4.6 -9.7 limestone

C215 222.5 -4.7 -10.0 limestone


C215 224.4∗∗ -5.2 -8.1 limestone


C215 224.9 -4.5 -8.3 limestone

C215 226.9 -4.9 -8.6 limestone

C215 228.1 -4.5 -9.5 limestone

C215 228.8 -4.5 -9.4 limestone

C215 229.3 -4.7 -9.0 limestone


C215 232.1 -4.7 -8.2 limestone

C215 233.3∗∗ -4.1 -9.4 limestone

C215 234.0 -4.1 -9.4 limestone

C215 234.5 -4.2 -9.6 limestone

C215 235.2 -4.2 -9.4 limestone


C215 237.5∗∗ -4.1 -9.9 limestone

C215 238.6 -3.9 -9.4 limestone

C215 239.7 -3.7 -9.3 limestone

C215 241.1∗∗ -3.7 -9.3 limestone

C215 242.3 -3.8 -10.0 limestone

C215 243.0 -3.7 -10.1 limestone


C215 245.0 -3.6 -10.0 limestone

C215 246.0 -3.6 -9.7 limestone

C215 247.2∗∗ -3.7 -10.2 limestone

C215 247.9 -3.5 -9.8 limestone


C215 249.9 -3.9 -9.8 limestone

C215 250.9 -3.4 -9.3 limestone

C215 251.9 -3.6 -10.4 limestone

C215 252.9 -3.6 -10.2 limestone


C215 255.0 -4.3 -10.5 limestone

27





C215 256.7 -3.1 -10.4 limestone

C215 257.5 -3.3 -10.0 limestone

C215 258.4 -3.1 -10.1 limestone


C215 261.0 -3.5 -10.8 limestone

C215 261.7∗∗ -2.9 -10.6 limestone

C215 262.7 -3.1 -10.7 limestone

C215 263.6 -3.5 -9.6 limestone


C215 265.2 -3.2 -10.6 limestone


C215 266.5 -3.3 -10.8 limestone

C215 266.8 -3.3 -10.9 limestone

C215 267.3 -3.3 -10.9 limestone

C215 268.4 -3.4 -11.0 limestone

C215 268.9 -3.6 -11.1 limestone

C215 269.4 -3.8 -12.2 limestone


C215 271.4 -3.4 -12.6 limestone

C215 271.6 -3.5 -12.7 limestone


C227 0.2 9.0 -9.6 limestone

C227 1.1 8.7 -9.9 limestone

C227 2.0 8.7 -9.8 limestone

C227 2.6 8.3 -9.9 limestone

C227 3.3 8.3 -10.6 limestone

C227 4.3 8.7 -8.5 limestone

C227 5.3 8.3 -11.3 limestone

C227 8.5 7.1 -9.9 limestone

C227 9.7††† 7.1 -10.3 -23.5 30.7 0.010 Rutgers limestone

C227 11.0∗∗ 8.1 -9.9 limestone

C227 12.9 6.9 -9.5 limestone

C227 14.0 6.6 -9.4 limestone

C227 15.0 7.6 -10.1 limestone

C227 15.6 7.2 -9.9 limestone

C227 17.0 7.1 -9.5 limestone

28




C227 18.5 7.9 -9.5 limestone

C227 20.0∗∗ 8.9 -6.5 limestone

C227 21.1∗∗ 8.4 -9.7 limestone

C227 22.2 7.5 -9.5 limestone

C227 23.5∗∗ 6.9 -7.9 limestone

C227 24.8 8.3 -9.8 limestone

C227 26.1∗∗ 8.0 -9.4 limestone

C227 27.3 7.9 -9.8 limestone

C227 28.6∗∗ 7.4 -11.1 limestone

C227 29.7∗∗ 6.2 -9.2 limestone

C227 30.8∗∗ 7.7 -9.3 limestone

C227 31.9 8.1 -9.7 limestone

C227 33.1 7.7 -9.8 limestone

C227 34.2∗∗ 8.2 -9.5 limestone

C227 35.4∗∗ 7.2 -9.6 limestone

C227 36.4†† 8.5 -9.6 -23.5 32.2 0.014 Rutgers limestone

C227 37.6 8.8 -9.3 limestone

C227 38.3 7.7 -9.8 limestone

C227 39.5 7.7 -9.8 limestone

C227 40.6∗∗ 7.0 -9.5 limestone

C227 41.8 6.4 -9.2 limestone

C227 43.0 6.7 -10.3 limestone

C227 44.6 7.3 -9.9 limestone

C227 45.6 8.0 -10.2 limestone

C227 46.7 8.2 -10.0 limestone

C227 47.7∗∗ 7.2 -7.2 limestone

C227 48.6 8.5 -10.6 limestone

C227 50.3 8.5 -10.6 limestone

C227 51.4 8.2 -10.6 limestone

C227 52.4††† 6.8 -10.9 -21.6 28.4 0.012 Rutgers limestone

C227 54.1 6.3 -10.3 limestone

C227 55.1 6.7 -11.6 limestone

C227 56.1∗∗ 5.9 -8.0 limestone

C227 57.8 6.2 -10.6 limestone

C227 58.7 6.9 -9.0 limestone

C227 60.7∗∗ 5.5 -10.7 limestone

C227 61.6 7.3 -11.1 limestone

29




C227 62.8 5.8 -11.0 limestone

C227 77.2 7.8 -10.8 limestone

C227 77.6∗∗ 7.3 -11.0 -24.2 31.5 0.021 Rutgers limestone

C227 85.2 6.0 -9.7 limestone

C227 86.4 9.8 -8.3 limestone

C227 87.7 9.4 -8.4 limestone

C227 89.1 9.8 -6.6 limestone

C227 90.3∗∗ 9.7 -8.2 limestone

C227 91.5 9.9 -9.1 limestone

C227 92.7 9.6 -10.4 -20.0 29.6 0.018 Rutgers limestone

C227 93.5 8.3 -10.3 limestone

C227 95.9 8.6 -10.4 limestone

C227 97.2 8.6 -10.4 limestone

C227 98.8 8.8 -10.2 limestone

C227 99.8 8.9 -9.2 limestone

C227 101.0 9.2 -9.4 limestone

C227 101.9∗∗ 7.9 -6.8 limestone

C227 103.0 8.1 -6.1 limestone

C227 104.0 8.7 -6.0 limestone

C227 105.1 9.9 -4.5 limestone

C227 106.4 9.8 -4.9 limestone

C227 107.8 9.7 -4.0 limestone

C227 108.8 9.6 -4.1 limestone

C227 109.9 9.7 -4.3 limestone

C227 110.0 9.5 -4.3 limestone

C227 112.1∗∗ 9.5 -4.6 limestone

C227 112.5 9.5 -4.9 limestone

C227 113.8 9.6 -4.9 limestone

C227 115.2 9.4 -5.6 limestone

C227 116.2∗∗ 8.7 -6.7 limestone

C227 117.4 9.2 -6.6 limestone

C227 118.6 9.3 -6.2 limestone

C227 119.6 9.3 -6.7 limestone

C227 120.9 9.2 -6.2 limestone

C227 121.9 9.4 -4.4 limestone


C227 123.9 9.5 -4.2 limestone

30




C227 125.0 9.6 -4.2 limestone

C227 126.2 9.6 -3.8 limestone

C227 127.2∗∗ 9.5 -4.1 limestone

C227 128.1 9.5 -5.0 limestone

C227 129.1 9.4 -4.4 limestone

C227 130.1 8.7 -4.8 limestone

C227 130.8 8.2 -5.5 limestone

C227 131.8 8.2 -6.9 limestone

C227 132.8 8.4 -7.2 limestone

C227 133.7 9.1 -3.3 limestone



C227 136.3 9.0 -4.7 limestone

C227 137.3 9.0 -4.2 limestone

C227 138.4∗∗ 9.0 -5.8 limestone

C227 139.8 7.8 -5.5 limestone

C227 141.4 7.5 -7.9 limestone

C227 142.4 8.4 -8.5 limestone

C227 143.0 8.3 -8.2 limestone

C227 144.2 8.6 -7.8 limestone

C227 145.2 8.5 -6.9 limestone

C227 146.3∗∗ 6.6 -8.5 limestone

C227 146.6∗∗∗ 0.2 -7.3 -26.0 26.2 0.017 Rutgers limestone

C227 147.6 8.0 -9.9 limestone

C227 149.6 7.9 -9.0 limestone

C227 150.1 8.3 -9.6 limestone

C227 153.8∗∗∗ 3.5 -5.5 limestone

C227 208.2∗∗∗ 5.4 -10.7 limestone

C227 212.8∗∗∗ 0.4 -9.2 limestone

C227 214.1 8.2 -7.0 limestone


C227 215.9 8.2 -6.5 limestone

C227 217.6 8.6 -6.4 limestone

C227 219.0∗∗ 8.2 -6.7 limestone

C227 220.4∗∗∗ 6.1 -5.2 limestone

C227 221.7∗∗ 5.6 -5.4 limestone


31




C227 223.9 8.3 -4.2 limestone

C227 225.1 8.9 -4.2 limestone

C227 226.4 8.1 -5.4 limestone

C227 227.6 8.9 -4.5 limestone

C227 228.7∗∗ 8.6 -3.7 limestone

C227 229.7 9.0 -3.8 limestone

C227 230.7∗∗ 4.5 -6.4 limestone

C227 231.5 8.3 -5.0 limestone

C227 232.7∗∗ 8.9 -4.4 limestone

C227 233.5 8.4 -4.5 limestone

C227 234.5∗∗ 9.0 -4.7 limestone

C227 235.5 9.0 -4.7 limestone

C227 236.6 8.1 -5.0 limestone

C227 237.6 8.3 -5.1 limestone


C227 240.0 8.8 -6.2 limestone

C227 240.9 8.6 -6.6 limestone

C227 246.0 7.8 -4.9 limestone

C227 247.3 9.0 -6.8 limestone

C227 248.2 8.8 -6.7 limestone

C227 249.1 9.2 -6.8 limestone

C227 250.3 9.3 -7.1 limestone

C227 251.5 9.0 -6.8 limestone

C227 252.5 9.3 -7.0 limestone

C227 253.8∗∗ 8.8 -6.2 limestone

C227 256.3 8.7 -6.3 limestone

C227 257.3 9.5 -7.2 limestone

C227 258.4 8.6 -6.1 limestone

C227 259.2 9.0 -6.7 limestone

C227 260.6 8.2 -5.6 limestone


C227 262.9∗∗ 9.0 -6.9 limestone

C227 264.1∗∗ 7.3 -7.0 limestone

C227 265.1∗∗ 5.6 -6.6 limestone

C227 266.1∗∗ 9.2 -6.4 limestone

C227 267.2∗∗ 8.3 -5.6 limestone


32




C227 269.2 7.1 -6.9 limestone

C227 270.4 7.5 -6.9 limestone

C227 271.3 7.5 -6.8 limestone

C227 272.4∗∗ 9.5 -6.4 limestone

C227 273.8 8.0 -5.1 limestone

C227 274.8 7.9 -6.5 limestone

C227 275.0 8.4 -6.3 limestone


C227 276.9∗∗ 9.0 -6.1 limestone

C227 277.9 8.5 -5.0 limestone

C227 278.5 9.3 -5.0 limestone

C227 278.7 9.4 -5.3 limestone

C227 279.8∗∗ 9.0 -6.8 limestone

C227 280.9∗∗ 8.3 -5.6 limestone

C227 281.9 8.5 -5.2 limestone

C227 282.9 8.7 -3.9 limestone

C227 283.9 9.1 -3.2 limestone

C227 285.0 8.6 -3.2 limestone

C227 289.1 8.4 -3.1 limestone

C227 290.2 8.7 -4.0 limestone

C227 292.6 8.3 -3.2 limestone

C227 298.2∗∗ 5.3 -4.4 limestone

C227 299.4∗∗ 1.3 -6.1 limestone

C227 301.2 8.4 -7.3 limestone


C227 305.6 8.8 -4.5 limestone

C227 306.6∗∗ 4.4 -5.6 limestone

C227 307.6 8.6 -4.9 limestone

C227 309.0 9.1 -4.8 limestone

C227 310.2∗∗ 9.2 -5.0 limestone

C227 311.3 9.4 -4.3 limestone

C227 312.4 9.1 -4.9 limestone


C227 314.5 9.2 -5.0 limestone

C227 315.8 9.5 -5.6 limestone

C227 316.8∗∗ 9.3 -5.3 limestone

C227 317.9 9.4 -5.9 limestone

33




C227 318.9 9.2 -5.5 limestone

C227 320.2∗∗∗ 6.3 -6.4 limestone

C227 321.2 9.7 -6.3 limestone

C227 323.3 9.6 -6.0 limestone

C227 324.1∗∗ 9.7 -6.2 limestone

C227 324.9 9.6 -7.2 limestone

C227 325.6 9.9 -7.8 limestone

C227 326.8 9.6 -8.1 limestone

C227 327.8 8.7 -7.7 limestone

C227 328.8 8.9 -8.6 limestone

C227 330.1∗∗∗ 1.3 -11.0 limestone

C227 331.5∗∗ 8.3 -8.8 limestone


C227 334.0 8.3 -9.5 limestone

C227 335.0∗∗∗ 6.1 -11.5 limestone

C227 396.2 7.7 -11.8 limestone

C227 397.0 8.3 -10.0 limestone

C227 398.6 9.4 -9.7 limestone

C227 400.0 9.4 -9.5 limestone

C227 401.0 9.8 -9.4 limestone


C227 403.4 9.7 -8.8 limestone

C227 405.3 8.9 -7.9 limestone

C227 406.4∗∗ 9.3 -8.7 limestone

C227 407.4 9.0 -8.3 limestone

C227 408.3 8.7 -8.9 limestone

C227 409.3 9.3 -8.6 limestone

C227 410.9 8.6 -8.3 limestone


C227 413.6∗∗ 8.6 -6.5 limestone

C227 417.1 9.8 -8.7 limestone

C227 418.1 9.8 -8.1 limestone

C227 419.1 9.5 -7.8 limestone

C227 420.2 -25.9 35.0 0.039 Rutgers limestone

C227 421.1 9.1 -7.8 limestone

C227 421.9∗∗ 8.4 -7.9 limestone

C227 422.9 8.8 -8.2 limestone

34




C227 423.9 8.9 -7.9 limestone

C227 424.4∗∗ 9.2 -8.4 limestone

C227 425.4 8.1 -8.3 limestone

C227 426.5∗∗ 8.9 -8.0 limestone

C227 429.5 9.0 -8.5 limestone

C227 430.3∗∗∗ 7.9 -7.7 limestone

C227 431.7∗∗ 7.8 -7.3 limestone

C227 432.7 7.2 -8.4 limestone

C227 433.6 6.9 -8.2 limestone

C227 434.7 8.8 -8.6 limestone

C227 435.7 8.7 -8.6 limestone

C227 436.5 8.9 -8.4 limestone

C227 437.6∗∗ 6.8 -6.3 limestone

C227 438.6∗∗ 7.7 -7.2 limestone

C227 439.5 8.6 -7.9 limestone

C227 440.3 8.4 -7.6 limestone

C227 440.4 9.7 -8.7 limestone


C227 442.7 9.3 -9.8 limestone

C227 443.7 8.7 -10.3 limestone

C227 444.3 8.8 -9.9 limestone

C227 445.3 9.2 -10.7 limestone

C227 448.0 6.8 -9.3 limestone

C227 451.9 7.7 -10.0 limestone

C227 452.9 8.4 -9.0 limestone

C227 453.9 8.9 -10.3 limestone

C227 454.9∗∗ 9.2 -9.2 limestone

C227 455.7 6.9 -10.3 limestone

C227 456.6 9.1 -10.2 limestone

C227 458.1 9.1 -9.5 limestone


C227 459.4 8.1 -10.8 limestone

C227 460.4 8.4 -11.1 limestone

C227 462.3 8.5 -11.4 limestone

C228 463.4††† -22.7 31.2 0.013 Rutgers limestone

C227 464.0 8.5 -11.5 limestone

C227 465.0∗∗ 8.7 -11.2 limestone

35




C227 465.8 8.1 -10.4 limestone


C227 468.2 8.0 -11.1 limestone

C227 469.2 8.1 -11.9 limestone

C227 470.2 8.0 -11.9 limestone

C227 471.3 7.5 -12.1 limestone

C227 472.3 7.9 -11.9 limestone

C227 473.3 6.7 -12.3 limestone

C227 701.0 8.4 -12.7 limestone


C227 703.1 9.2 -12.3 limestone

C227 704.1∗∗ -1.1 -8.6 limestone

C227 704.9 5.4 -10.5 limestone

C227 711.5∗∗ 5.3 -9.8 limestone

C227 712.6 2.3 -9.5 limestone

C227 713.6∗∗ -2.0 -7.7 limestone

C227 714.5 8.1 -9.6 limestone

C227 713.8 3.0 -8.4 limestone

C227 716.8 9.0 -9.2 limestone


C227 718.6 8.7 -9.1 limestone

C227 719.7∗∗ -2.3 -7.6 limestone

C227 720.9 5.9 -8.8 limestone

C227 721.9 9.2 -9.0 limestone

C227 723.0 9.5 -7.9 limestone

C227 724†† 9.1 -10.2 -24.6 33.7 0.008 Rutgers limestone

C227 724.9∗∗ 0.6 -8.1 limestone

C227 725.8 9.0 -10.4 limestone

C227 727.0∗∗ 4.4 -9.2 limestone


C227 729.3 8.2 -10.4 limestone

C227 730.3 8.3 -11.0 limestone


C227 732.4 7.3 -9.6 limestone

C227 733.5 7.8 -11.2 limestone


C227 735.5 7.1 -11.5 limestone

36




C227 736.5∗∗ 6.6 -11.7 limestone


C227 782.3 7.2 -9.5 limestone

C227 783.4 7.8 -11.7 limestone

C227 784.1 7.9 -11.4 limestone

C227 795.9 9.1 -11.9 limestone

C227 796.9 9.4 -11.9 limestone

C227 797.9 9.4 -11.3 limestone


C227 800.3 8.9 -10.7 limestone

C227 801.4 9.2 -11.0 limestone

C227 802.4 9.3 -11.0 limestone

C227 832.9 5.1 -10.9 limestone

C227 834.1∗∗ 6.5 -11.9 limestone

C227 835.2 7.1 -12.2 limestone


C227 838.4 8.7 -12.0 limestone

C227 839.4 8.7 -11.9 limestone

C227 840.5 9.0 -11.5 limestone

C227 841.5 8.7 -11.5 limestone

C227 842.6 9.1 -11.3 limestone

C227 843.9 9.0 -11.4 limestone

C227 845.1 8.9 -11.1 limestone

C227 846.1 8.7 -11.7 limestone

C227 847.0 7.9 -10.9 limestone

C227 849.5 8.1 -12.3 limestone

C227 850.5††∗∗ 8.4 -12.5 -25.0 33.4 0.008 Rutgers limestone

C227 851.6 8.3 -12.2 limestone

C227 853.6 7.4 -12.8 limestone

C227 854.7 8.1 -14.2 limestone

C227 857.1 7.9 -11.5 limestone

C227 858.2 8.1 -12.0 limestone

C227 859.2 7.7 -11.4 limestone

C227 860.2 7.8 -11.6 limestone

C227 861.2 7.5 -11.7 limestone


C227 864.6 8.0 -12.1 limestone

37




C227 868.7 7.9 -12.5 limestone

C227 869.7 7.8 -12.8 limestone

C227 870.6 7.5 -12.6 limestone


C227 872.4 6.5 -12.9 limestone

C227 873.4 6.0 -12.9 limestone

C227 874.5 6.1 -13.0 limestone

C227 876.0∗∗ 5.9 -13.2 limestone

Notes: All δ13Corg values were obtained from analysis of insoluble residues isolated through decarbonation of limestone or dolostone

samples. ∗ symbols indicate that multiple δ13Ccarb analyses have been averaged into the value presented with the number of ∗ corre-

sponding to the number of analyses of the sample. The † symbols represent the same but for δ13Corg analyses. ∆δ13C was determined

through subtraction of δ13Corg from δ13Ccarb. This approximation of the difference between the two values avoids confusion with the

primary εp that is associated with the fraction between DIC and primary biomass as they are not one and the same. The Corg lab column

indicates whether the organic carbon data was collected at the University of Adelaide or at Rutgers University. All δ13Ccarb data were

generated at the University of Michigan.

38

cryogenian glaciation and the onset of carbon-isotope decoupling … · 2013. 7. 30. · values...

Documents