Natural radionuclide mobility and its influence on U–Th–Pb dating of secondary minerals from the unsaturated zone at Yucca Mountain, Nevada

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Geochimica et Cosmochimica Acta 72 (2008) 2067–2089 www.elsevier.com/locate/gca

Natural radionuclide mobility and its influence on U–Th–Pb dating of secondary minerals from the unsaturated zone at Yucca Mountain, Nevada L.A. Neymark a,*, Y.V. Amelin b,c a U.S. Geological Survey, DFC Box 25046, MS 963, Denver, CO 80225, USA Geological Survey of Canada, 601 Booth Street, Room 693, Ottawa, Ont., Canada K1A 0E8 c Research School of Earth Sciences and Planetary Science Institute, The Australian National University, Canberra, ACT 0200, Australia b

Received 21 June 2007; accepted in revised form 6 February 2008; available online 16 February 2008

Abstract Extreme U and Pb isotope variations produced by disequilibrium in decay chains of 238U and 232Th are found in calcite, opal/chalcedony, and Mn-oxides occurring as secondary mineral coatings in the unsaturated zone at Yucca Mountain, Nevada. These very slowly growing minerals (mm my1) contain excess 206Pb and 208Pb formed from excesses of intermediate daughter isotopes and cannot be used as reliable 206Pb/238U geochronometers. The presence of excess intermediate daughter isotopes does not appreciably affect 207Pb/235U ages of U-enriched opal/chalcedony, which are interpreted as mineral formation ages. Opal and calcite from outer (younger) portions of coatings have 230Th/U ages from 94.6 ± 3.7 to 361.3 ± 9.8 ka and initial 234 U/238U activity ratios (AR) from 4.351 ± 0.070 to 7.02 ± 0.12, which indicate 234U enrichment from percolating water. Present-day 234U/238U AR is 1 in opal/chalcedony from older portions of the coatings. The 207Pb/235U ages of opal/chalcedony samples range from 0.1329 ± 0.0080 to 9.10 ± 0.21 Ma, increase with microstratigraphic depth, and define slow longterm average growth rates of about 1.2–2.0 mm my1, in good agreement with previous results. Measured 234U/238U AR in Mn-oxides, which pre-date the oldest calcite and opal/chalcedony, range from 0.939 ± 0.006 to 2.091 ± 0.006 and are >1 in most samples. The range of 87Sr/86Sr ratios (0.71156–0.71280) in Mn-oxides overlaps that in the late calcite. These data indicate that Mn-oxides exchange U and Sr with percolating water and cannot be used as a reliable dating tool. In the U-poor calcite samples, measured 206Pb/207Pb ratios have a wide range, do not correlate with Ba concentration as would be expected if excess Ra was present, and reach a value of about 1400, the highest ever reported for natural Pb. Calcite intergrown with opal contains excesses of both 206Pb and 207Pb derived from Rn diffusion and from direct a-recoil from U-rich opal. Calcite from coatings devoid of opal/chalcedony contains 206Pb and 208Pb excesses, but no appreciable 207Pb excesses. Observed Pb isotope anomalies in calcite are explained by Rn-produced excess Pb. The Rn emanation may strongly affect 206Pb–238U ages of slow-growing U-poor calcite, but should be negligible for dating fast-growing U-enriched speleothem calcite. Published by Elsevier Ltd.

1. INTRODUCTION Records of ancient water flow in arid environments are preserved in precipitated from ground water mineral coatings, composed largely of calcite (CaCO3), opal (SiO2 *

Corresponding author. Fax: +1 303 236 4930. E-mail address: [email protected] (L.A. Neymark).

0016-7037/$ - see front matter Published by Elsevier Ltd. doi:10.1016/j.gca.2008.02.001

nH2O), chalcedony (cryptocrystalline quartz), and Fe–Mn oxides. Sedimentary carbonates, such as corals, tufas, and speleothems, can be dated precisely by radiocarbon, and U-series techniques, but are limited to samples younger than ca. 40 and 500 ka, respectively. U–Pb method allows direct dating of older speleothems (Richards and Dorale, 2003; Woodhead et al., 2006), making it possible to correlate environmental and geologic events preserved in these

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terrestrial records. The U–Pb system in opal and chalcedony allows dating in the age range from 50 ka to millions of years and older (Ludwig et al., 1980; Neymark et al., 2000, 2002). Recently, the reliability of U–Pb dating of opal was questioned on the basis of idea that large Rn-produced excesses of 206Pb and 207Pb would result in both 206Pb/238U and 207Pb/235U ratios that grossly overestimate mineral deposition ages (Pashenko and Dublyansky, 2002a,b, 2006; Dublyansky et al., 2003). These authors conclude that Rn emanation renders the U–Pb dating of young hydrogenic minerals formed in open cavities unreliable, and suggest that U–Pb dating of Quaternary speleothems is also problematic (Dublyansky and Pashenko, 2004). Secondary coatings of calcite, opal/chalcedony, and Mn-oxides exist in fractures and lithophysal cavities (primary voids formed during cooling) in felsic tuffs at Yucca Mountain, Nevada. The site is proposed to host a geologic repository of high-level nuclear waste. Meteoric water plays a crucial role in transporting radionuclides from the proposed repository and hydrogenic minerals record a history of water movement through the approximately 700-meter (m)-thick unsaturated zone (UZ). The timing of the secondary mineral deposition has been established by 230Th/U and U–Pb dating of U-enriched opal and chalcedony (Neymark and Paces, 2000; Neymark et al., 2000, 2002; Paces et al., 2001, 2004; Wilson et al., 2003; Nemchin et al., 2006). Isotopic ages indicated that secondary minerals started to form in the Yucca Mountain UZ shortly after the eruption of the host tuffs 12.7 to 12.8 million years ago (Ma) (Sawyer et al., 1994; Huysken et al., 2001) and that deposition continued at very slow long-term average rates from <1 to about 5 millimeters per million years (mm my1) throughout the history of the mountain (Neymark et al., 2002; Paces et al., 2004). Calcite is about 10 times more abundant than opal/chalcedony in the Yucca Mountain UZ (Paces et al., 2001) and an evaluation of calcite as a viable U–Pb geochronometer is important for its potential usage in refining the history of water movement through the Yucca Mountain UZ. Slow mineral growth rates based on the opal/chalcedony U-series and U–Pb ages were questioned and it was proposed instead that the calcite–silica deposits may have formed rapidly and possibly recently from upwelling hydrothermal water (Dublyansky et al., 2001, 2003, 2005). In order to reinterpret published U–Pb ages, a ‘‘radon emanation” model was proposed (Pashenko and Dublyansky, 2002a,b, 2006; Dublyansky et al., 2003), which hypothesizes that cavities were filled with water that contained 222 Rn and 219Rn derived from U decay in the surrounding tuffs. In this study, we test the possible presence and mobility of excess Rn and other intermediate decay products of U in secondary minerals from the Yucca Mountain, Nevada using Sr, U, Th, and Pb isotopic compositions and selected element concentrations in secondary calcite, opal/chalcedony, and Mn-oxides. Calcite without admixed opal from the Yucca Mountain UZ contains little U (typically <0.05 lg g1, Paces et al., 2001) and has much lower 238 U/204Pb ratios than opal/chalcedony (Neymark et al., 2003). Therefore, calcite is a good target for determining

the presence of U-unsupported excess Pb isotopes, because contributions of radiogenic 206Pb and 207Pb from in situ U decay should be much smaller than in coexisting opal/chalcedony. U–Pb isotope measurements of these secondary minerals address their potential to affect the mobility for U, Th, and decay products and as U–Th–Pb mineral geochronometers. The origin of radiogenic Pb unsupported by in situ U decay and its influence on calculated 206 Pb/238U and 207Pb/235U ages of calcite and opal/chalcedony are also evaluated. 2. GEOLOGICAL SETTING AND SAMPLES Yucca Mountain in southwestern Nevada (Fig. 1A) consists of a 1–3 km-thick sequence of 12.8–12.7 Ma (Sawyer et al., 1994; Huysken et al., 2001) felsic welded and nonwelded tuffs of the Paintbrush Group of Miocene age, 500–700 m of which is located above the modern water table. Samples for this study were collected in the Exploratory Studies Facility (ESF), a 7.6-m-diameter, 7.8-kmlong, C-shaped tunnel (Fig. 1A), which was excavated into Yucca Mountain through the welded part of the Tiva Canyon Tuff (TCw), non-welded Paintbrush Tuff (PTn), and welded part of the Topopah Spring Tuff (TSw) (Fig. 1B). Calcite and silica deposits in the ESF are found predominantly in open fractures and lithophysal cavities (Paces et al., 2001; Whelan et al., 2002, 2004; Wilson et al., 2003; Wlson and Cline, 2005). A generalized paragenetic sequence presented by Whelan et al. (2002) indicates initial high-temperature deposition of vapor-phase tridymite, cristobalite, quartz, feldspar, hematite, zeolites, and clays along early cooling joints and fractures, and lithophysal cavity rims. Vapor-phase minerals were followed by deposition of Mn-oxide minerals, early stage calcite, fluorite, and chalcedony–quartz, and then intermediate- and late-stage calcite with minor amounts of opal (Whelan et al., 2002). The complex growth sequences of a late-stage calcite blade containing opal are illustrated by successively magnified scanning electron microscope (SEM) images (Fig. 2). X-ray diffraction (XRD) data for the secondary silica showed a systematic change with microstratigraphic depth in coatings so that small amounts of cryptocrystalline quartz were observed only in subsamples with 207Pb/235U ages older than 4 Ma (Neymark et al., 2003). The crystallinity of the quartz in the silica subsamples increased with age and was interpreted as slow maturation (crystallization) of originally amorphous silica. Because opal and chalcedony share many geochemical characteristics, they are denoted as opal/chalcedony in this paper when those characteristics are discussed. All samples of Mn-oxides were attached to unaltered fracture surfaces or to vapor-phase minerals on floors of lithophysal cavities and were always stratigraphically older than the oldest portions of calcite–silica coatings present in the same cavities (Fig. 3A). The XRD analyses (Carlos et al., 1993; Vaniman and Chipera, 1996) showed that Mn-oxides in the Yucca Mountain UZ are represented mainly by rancieite ½ðCa; Mn2þ ÞMn4þ 4 O9  3H2 O and lithiophorite ½Li2 Al8 ðMn2þ ; Co; NiÞMn4þ 10 O35  14H2 O. Mn-oxides commonly form small crusts and dendrites

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals 116º30’

2069

116º25’

36º55’

NEVADA

Y uc

Carson City

ca M

Approximate area of map

ESF

ount

36º50’

Las Vegas

ain

Quaternary alluvium Miocene felsic tuffs Pliocene and

36º45’

Quaternary basalts Major faults Exploratory Studies

95

36º40’

TSw TCw

PTn

TCw

Facility (ESF)

TSw

1160 High-Silica Rhyolite Quartz Latite

Elevation (m)

1140

1120

1100 1080

1060

0

2000

4000

6000

8000

ESF Distance (m) Fig. 1. (A) Map showing Yucca Mountain and the Exploratory Studies Facility (ESF) tunnel (after Scott and Bonk, 1984). (B) Sampling sites in the ESF tunnel shown relative to the elevation. Hydrogeologic units non-welded Paintbrush Tuff (PTn), welded Topopah Spring Tuff (TSw), and welded Tiva Canyon Tuff (TCw) and types of rocks (quartz latite and high-silica rhyolite) also are shown.

(sometimes on smectite) and despite the small volume of these minerals they may be effective in water–rock interactions because of very large specific surface areas (Fig. 3). Secondary mineral coatings were collected from 20 sampling sites in the ESF (Fig. 1B, Table EA1 in Electronic Annex 1 and Table 1). The term ‘‘sample” is used in this paper for a hand specimen of secondary mineral coating, with or without its host tuff, collected from a particular fracture or cavity. A sample can comprise one or more mineral species with one or more occurrences of mineral species. If several occurrences of a mineral are found within a sample, their composition, and presumably age, may vary. A part of a sample representing a single occurrence of one mineral, extracted for analysis, is called a subsample. In some cases subsamples were divided into still smaller parts, which were analyzed separately to test homogeneity of subsamples.

3. ANALYTICAL TECHNIQUES 3.1. ICP-MS chemical analyses Concentrations of selected elements in secondary mineral samples were determined by PlasmaQuad-IIITM inductively coupled plasma mass spectrometer (ICP-MS) at the U.S. Geological Survey (USGS) Yucca Mountain Project Branch laboratory in Denver. Inclusion-free calcite fragments were hand-picked under a binocular microscope and most of the fragments were completely digested at room temperature overnight in a 1% nitric acid (HNO3) solution. In cases where residue was still present it was digested on a hot plate in concentrated hydrofluoric acid (HF) and combined with the HNO3 digestion. A set of ICP-MS grade element standards was used for the instrument calibration and precisions of the reported concentra-

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Fig. 2. Scanning electron microscope (SEM) images at successively greater magnification of a single blade of calcite (Cc) and opal (Op) from a 17-mm-thick coating in a cavity, sample HD2057, Exploratory Studies Facility (ESF) tunnel, distance 2962.2 m from the north portal, Yucca Mountain, Nevada (Paces et al., 2001).

tions are about 10%.1 REE concentrations in Cl chondrites used to normalize measured values are from Taylor and McLennan (1985, p. 298). The 206Pb/207Pb ratios analyzed during ICP-MS runs of seven calcite subsamples from the coating HD2089A have 2% uncertainties. A whole-rock sample of host high-silica rhyolite was analyzed by the U.S. Geological Survey’s Rock Analysis Laboratory using methods described by Arbogast (1996). 3.2. TIMS Sr, U–Pb, and U-series isotope analyses Isotopic compositions of Sr in Mn-oxide subsamples were determined at the USGS Yucca Mountain Project Branch laboratory in Denver. Powdered minerals collected using dental burs were dissolved in 7 N HNO3, Sr was separated on an ion-exchange column, and isotope ratios were measured on a thermal ionization mass-spectrometer (TIMS) Finnigan MAT-262TM in static mode. Total blank of less than 1 nanogram (ng) Sr is negligible compared to several hundred ng Sr in the analyzed samples. 87Sr/86Sr 1

All errors cited in this paper are ±2r unless stated otherwise.

ratios are adjusted to the value of 0.70920 in the modern seawater Sr standard EN-1 (Ludwig et al., 1988). Total uncertainty for 87Sr/86Sr values including external error is ±0.00005. U–Pb and 230Th–U analyses of calcite and opal/chalcedony were done at the Jack Satterly Geochronology Laboratory, Royal Ontario Museum (ROM), Toronto, Canada, using techniques described by Neymark et al. (2000, 2002). The hand-picked inclusion-free mineral fragments were spiked with a mixed 205Pb–229Th–233U–236U tracer solution before dissolution, and U–Pb and U-series isotopic data were obtained from the same aliquot. Isotopic compositions of U, Th, and Pb were measured with a TIMS VG354TM, equipped with ion-counting Daly photomultiplier detectors. The total procedural blanks measured at ROM during this study ranged from 0.3 to 0.9 picogram (pg) Pb and from 0.3 to 2.0 pg U and Th. Isotopic compositions of U, Th, and Pb in Mn-oxides were determined at the USGS Yucca Mountain Project Branch laboratory in Denver. U and Pb isotope analyses were conducted on separate splits of the bulk mineral powders analyzed for Sr isotopes after dissolution in 7 N HNO3 and separation on an ion-exchange column. Additional

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

2071

Fig. 3. Samples of secondary mineral coatings from the Exploratory Studies Facility (ESF), Yucca Mountain, Nevada. (A) Photograph of sample HD2089A at distance 3991.4 m from the north portal, showing relations between the vapor-phase minerals (VPh), Mn-oxides, and calcite in a coating from the floor of the lithophysal cavity. The tip of an analyzed calcite blade is shown in the insert. (B, C, and D) Scanning electron microscope (SEM) images at successively greater magnification showing a surface of the Mn-oxide crust shown in (A). (E and F) SEM images of sample HD2062, ESF distance 3107.4 m from the north portal showing two different Mn-oxides on a fracture surface: lithiophorite (E) and rancieite (F).

Mn-oxide samples (with Mn2 in the subsample name; Tables 1 and 2) were separated from the same coatings for U and Th isotope analyses. The isotopic analyses were conducted using Finnigan MAT 262TM and ThermoFinnigan TritonTM TIMS equipped with ion counting systems

and an energy filter (TritonTM). Isotope ratios were measured in a static mode using Faraday cups for Pb and a peak jumping mode using an ion counter for U and Th. U, Th, and Pb blank contributions were negligible during the Mn-oxide analyses. Measured Pb-isotope ratios were

2072

Subsample name

204

206

Pb/204Pbb Error (%)

ESF Occurrence Mineral distance a (m)

Coating Sample thickness weight (mm) (mg)

U (lg/g)

HD1939Pb1-Cc HD1939-Mn2

818.2 ”

Fracture ”

Calcite Mn-oxide

3.5 <0.5

1.08E01 1.36E01 5.05E+01 1.83E+00 1.847E+01 8.19E+00 —c 1.52E+05 — —

HD2055Pb6-Cc

2911.2

Cavity

Calcite

20.0

HD2055Pb7-Cc1 HD2055Pb7-Cc2 HD2055Pb10-Cc HD2055Pb11-Cc HD2055Pb11-Op HD2055Pb12-Cc HD2055Pb12-Op

” ” ” ” ” ” ”

” ” ” ” ” ” ”

” ” ” ” Opal Calcite Opal

” ” ” ” ” ” ”

13.16 1.80 3.60 33.10 5.24 30.98 1.19

1.07E02 3.71 E02 1.36E02 1.11E01 2.00E02 9.27E02 3.62E02 2.04E02 1.00E01 7.31E03 1.86E+02 2.85E02 1.02E01 1.41E02 2.45E+02 1.96E01

HD-2057-Pb1-Cc HD-2057-Pb2-Cc HD-2057-Pb2-Op

2962.2 ” ”

Cavity ” ”

Calcite ” Opal

17.0 ” ”

8.6 12.61 7.63

HD2059Pb4-Cc

3017.78 Cavity

Calcite

18.3

45.05

HD2062Pb1-Cc HD2062Pb2-Cc HD2062Pb3-Mn HD2062-Mn2

3107.4 ” ” ”

Calcite ” Mn-oxide ”

3.6

109.89 35.42 0.698 4.060

Fracture ” ” ”

<0.5 <0.5

38.91 1.88 1.466

Pb (lg/g) Th (ng/g)

Pb (ng/g)

207

Pb/204Pbb Error (%)

208 204

Pb/ Pbb

Error (%)

0.09 —

1.560E+01 —

0.14 —

39.24 —

0.18 —

206 207

—

Pb/ Pbb

Th/U

238

U/204Pbb

232

Th/204Pbb

1.184 4.68E+02 5.01E+01 1.86E+04 —

2.29E+04 —

1.327 —

1.84E+01

—

—

4.97E01 2.061E+01

0.30

1.554E+01

0.34

37.53

0.40

— 2.20E+00 — — 3.29E01 — 4.22E+01

1.48E+00 1.24E+00 2.69E01 3.04E02 1.37E02 3.78E02 3.02E02

0.092 0.22 0.16 0.53 2.2 0.47 3..2

1.559E+01 1.562E+01 1.561E+01 1.619E+01 3.607E+01 1.642E+01 2.500E+02

0.14 0.24 0.20 0.23 1.8 0.21 3.1

39.18 39.28 39.30 39.97 41.28 39.91 44.51

0.94 0.32 0.97 1.6 1.7 0.25 2.2

1.217 1.227 1.285 11.35 55.56 19.25 24.83

7.80E+01 1.38E+01 1.14E+02 2.81E+03 1.16E+07 2.30E+03 6.89E+06

— 1.48E+0.3 — — 1.99E+04 — 1.16E06

2.23E02 8.83E03 1.12E01 5.36E02 1.094E+02 1.09E01 1.54E02 3.20E01 1.60E01 4.135E+01 9.56E+01 1.97E02 2.15E01 1.30E02 1.449E+03

0.67 0.20 3.62

1.559E+01 1.557E+01 3.347E+01

0.22 0.2 2.03

39.20 38.43 39.72

0.244 0.229 0.632

7.016 5.00E+00 3.60E+02 2.655 2.93E+00 5.89E+02 43.31 2.25E03 6.35E+06

1.76E+03 1.69E+03 1.39E+04

2.35E02 1.35E02 4.85E02 9.71E02 8.259E+01

0.22

1.585E+01

0.25

39.52

0.30

5.211 2.07E+00 2.06E+02

4.15E+02

0.50 0.12 0.058 —

1.572E+01 1.558E+01 1.560E+01 —

0.21 44.83 0.17 39.14 0.087 39.18 — —

0.30 0.18 0.12 —

6.85E02 2.90E02 1.66E+02 1.00E+02

4.78E03 3.46E02 2.60E+03 —

— 1.55E+01 4.49E+04 1.19E+05

1.02E02 4.62E01 3.51E+04 —

1.898E+01 1.916E+01 2.005E+01 1.838E+02 2.004E+03 3.162E+02 6.205E+03

4.074E+02 1.918E+01 1.816E+01 —

25.92 1.231 1.164 —

— 1.10E+02 — — 1.77E03 — 1.72E01

— 5.34E+02 2.71E+02 1.18E+03

5.72E+01 5.34E+01 4.01E+00 —

HD2065Pb4-Cc

3316.2

Cavity

Calcite

18.0

25.33

3.31E02 5.90E03 2.21E01 4.83E02 6.644E+01

0.19

1.567E+01

0.23

39.18

0.25

HD2074Pb1-Cc2 HD2074Pb1-Cc3 HD2074Pb2-Cc1 HD2074Pb2-Cc2 HD2074Pb2-Op

3050.7 ” ” ” ”

Cavity ” ” ” ”

Calcite ” ” ” Opal

36.7 ” ” ” ”

66.14 60.48 89.88 40.26 0.75

4.29E02 4.08E02 3.52E02 3.84E02 1.40E+02

5.17E03 2.56E03 1.42E03 1.77E03 1.23E02

— 2.84E02 — 1.08E02 1.29E+00

8.41E03 2.85E02 9.80E03 1.86E02 3.08E02

5.550E+02 3.384E+01 8.882E+01 3.999E+01 3.423E+02

1.0 0.29 0.56 0.38 4.7

1.553E+01 1.561E+01 1.585E+01 1.573E+01 1.934E+01

0.33 0.33 0.29 0.43 1.6

43.61 39.18 39.48 38.61 38.05

0.77 0.42 0.33 0.43 1.8

35.74 2.168 5.603 2.543 17.70

— 6.95E01 — 2.81E03 9.19E03

4.34E+03 1.22E+03 3.05E+03 1.76E+03 3.87E+06

— 8.25E+02 — 4.82E+02 3.47E+04

HD2089APbl-Ccl HD2089APb1-Cc2a HD2089APb1-Cc2b HD2089APb1-Cc2c HD2089APb1-Cc3 HD2089APb2-Cc HD2089APb3-Mn

3991.4 ” ” ” ” ” ”

Cavity ” ” ” ” ” ”

3.2 ” ” ” ” ” <0.5

48.48 61.22 55.14 11.64 2.33 26.23 5.49

8.64E03 1.50E02 3.30E02 1.90E02 4.07E02 8.06E03 1.03E+01

1.87E02 2.22E02 3.30E02 2.84E02 3.72E02 1.37E02 2.01E+02

— — — — 1.61E01 8.34E+00 3.43E+04

6.75E03 1.61E03 1.36E03 1.37E02 3.06E02 1.23E01 2.71E+03

2.686E+03 1.349E+04 2.385E+04 1.991E+03 1.145E+03 5.557E+01 1.860E+01

1.8 6.3 8.5 3.5 1.6 0.14 0.060

1.572E+01 1.535E+01 1.675E+01 1.567E+01 1.560E+01 1.558E+01 1.560E+01

0.60 72.02 2.9 210.2 4.5 341.1 0.57 65.93 0.71 53.63 0.17 39.62 0.090 39.17

1.0 5.4 7.9 1.6 0.80 0.20 0.12

170.9 879.3 1423 127.1 73.43 3.568 1.192

— — — — 3.96E+00 1.04E+03 2.71E+03

1.09E+03 7.91E+03 1.07E+04 1.18E+03 1.13E+03 5.58E+00 3.24E+00

— — — — 4.36E+04 5.63E+04 1.05E+04

HD2089A-Mn2

”

”

Calcite ” ” ” ” ” Fe-Mnoxide ”

”

6.50

1.72E+01 —

7.63E+04 —

—

—

—

—

—

—

4.241 6.68E+00 5.83E+02

— 2.78E+04 1.60E+03 —

—

4.44E+03 —

3.79E+03

—

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

Table 1 U, Th, and Pb concentrations, and Pb isotope compositions in samples of secondary minerals from the Exploratory Studies Facility (ESF), unsaturated zone, Yucca Mountain, Nevada

HD2092Pb1-Cc HD2092Pb1-Mn HD2092-Mn2

754d ” ”

” ” ”

Calcite Mn-oxide ”

1.3 <0.5 ”

40.66 0.36 7.83

9.80E02 2.36E02 1.08E+02 3.10E01 1.975E+01 6.86E+00 1.82E+03 1.47E+05 2.47E+04 1.817E+01 9.26E+00 — 2.73E+05 — —

0.14 1.561E+01 0.060 1.558E+01 — —

0.18 0.090 —

39.69 0.22 39.054 0.12 — —

1.266 1.0E+03 2.69E+02 1.166 2.14E+04 2.37E01 — 2.94E+04 —

2.88E+05 4.94E+03 —

HD2098Pb3-Cc

2678.1

”

Calcite

1.3

33.47

4.15E03 2.26E02 5.07E+00 2.79E01 2.495E+01

0.11

1.558E+01

0.15

39.33

0.20

1.601 1.22E+03 2.26E+01

1.51E+04

HD2109Pb1-Cc

4714

”

Calcite

1.8

60.50

4.07E03 3.17E02 2.00E+01 3.98E01 2.353E+01

0.13

1.562E+01

0.17

39.48

0.21

1.506 4.91E+03 8.70E+00

4.17E+04

117.53

5592.8

Fracture

Calcite

3.4

6272.06 ” ” ”

” ” ” ”

Calcite ” Mn-oxide ”

1.2 ” <0.5 ”

53.12 53.10 3.87 2.56

5.00E03 4.70E02 2.01E+00 1.50E+01

1.65E02 4.95E03 9.70E01 5.43E02 3.494E+01

HD2227Pb1-Cc HD2227Pb1-Mn HD2227-Mn2

4117.6 ” ”

Cavity ” ”

Calcite Mn-oxide ”

0.8 <0.5 ”

35.50 2.53 6.06

6.50E02 1.90E02 1.92E+01 2.22E01 2.927E+01 1.24E+01 3.07E+02 8.11E+04 4.16E+03 1.817E+01 1.74E+01 — 9.66E+04 — —

1.38E01 6.49E03 7.19E02 —

8.19E+00 1.33E+00 2.15E+04 1.95E+05

5.58E01 8.22E02 9.72E+02 —

1.907E+02 2.325E+01 1.818E+01 —

0.17

1.559E+01

0.21

39.61

0.23

0.15 0.26 0.060 —

1.560E+01 1.561E+01 1.559E+02 —

0.18 0.31 0.090 —

40.17 39.16 39.16 —

0.22 0.43 0.12 —

12.22 1.490 1.166 —

2.242 5.86E+01 1.559E+010.2 1.48E+04 2.59E+02 4.87E+02 1.76E+00 —

1.22E+04 1.35E+04 1.83E+04 —

0.1 0.06 —

1.558E+01 1.557E+01 —

0.16 0.090 —

39.55 39.09 —

0.21 0.12 —

1.878 2.95E+02 2.49E+02 1.167 6.54E+03 1.54E+00 — 5.57E+03 —

7.16E+04 1.62E+04 —

1.64E+03 2.84E+01 1.07E+04 1.30E+04

HD2231Pb1-Cc

6784.0

”

Calcite

3.4

22.58

2.51E+00 2.19E01 1.24E+03 2.95E+00 1.852E+01

0.30

1.565E+01

0.33

39.18

0.35

7.24E+02

3.49E+05

HD2233Pb1-Ch1 HD2233Pb1-Ch2 HD2233Pb2-Ch1 HD2233Pb2-Ch2 HD2233Pb3-Ch HD2233Pb4-Ch HD2233Pb5-Ch

6838.1 ” ” ” ” ” ”

” ” ” ” ” ” ”

Chalcedony ” ” ” ” ” ”

10.0 ” ” ” ” ” ”

2.32 3.75 3.51 4.52 3.33 3.39 2.92

1.28E+01 1.70E+01 2.14E+01 2.01E+01 1.68E+01 4.41E+00 3.84E+00

2.43E02 2.32E02 3.27E02 2.72E02 3.44E02 9.78E03 1.92E02

— 12.2E02 — 3.19E02 — — —

6.48E02 6.34E02 1.56E01 1.05E01 2.08E01 4.57E02 1.73E01

3.032E+02 2.941E+02 1.469E+02 1.933E+02 1.047E+02 1.515E+02 5.365E+01

0.92 0.88 0.33 0.25 0.25 0.95 0.36

2.075E+01 2.244E+01 1.986E+01 2.151E+01 1.849E+01 2.105E+01 1.685E+01

0.43 0.80 0.31 0.21 0.23 0.62 0.34

50.04 48.89 41.76 43.59 40.93 40.56 39.50

0.47 0.79 2.1 0.25 0.27 4.8 0.46

14.61 13.10 7.394 8.983 5.661 7.200 3.184

— 7.20E04 — 1.59E03 — — —

1.68E+05 2.28E+05 1.17E+05 1.63E+05 6.85E+04 8.21E+04 1.89E+04

— 1.60E+02 — 2.52E+02 — — —

HD2247Pb1-Cc1 HD2247Pb1-Cc2 HD2247Pb2-Cc HD2247Pb1-Op HD2247Pb1-Mn HD2247-Mn2

7165.8 ” ” ” ” ”

” ” ” ” ” ”

Calcite ” ” Opal Mn-oxide ”

1.1 ” ” ” <0.5 ”

37.12 45.87 77.39 23.27 1.78 22.70

3.29E02 6.74E02 2.83E02 2.60E+02 8.24E+00 3.70E+00

3.70E02 3.78E02 1.75E02 1.95E01 7.32E+02 —

5.94E+00 1.84E+01 1.30E+00 2.51E+02 3.98E+04 2.10E+04

4.94E01 5.06E01 2.29E01 9.90E02 9.91E+03 —

1.911E+01 1.891E+01 2.057E+01 8.872E+01 1.813E+01 —

0.11 1.11 0.17 0.11 0.06 —

1.559E+01 1.559E+01 1.559E+01 1.829E+01 1.559E+01 —

0.15 0.15 0.21 0.15 0.090 —

31.22 39.24 39.29 38.75 39.14 —

0.23 0.19 0.23 0.192 0.12 —

1.226 1.213 1.319 4.850 1.163 —

1.80E+02 2.73E+02 4.60E+01 9.64E01 4.83E+03 5.69E+03

5.67E+01 1.13E+02 1.69E+05 7.08E01 7.08E01 —

9.96E+03 3.02E+04 4.73E+03 1.58E+05 3.33E+03 —

HD2293Pb1-Cc HD2293Pb1-Op

103.9 ”

” ”

Calcite Opal

5.0 ”

102.48 0.97

6.95E02 2.29E02 — 2.47E01 3.671E+01 5.03E+01 4.24E02 1.76E+00 4.37E02 8.873E+02

0.10 2.7

1.573E+01 3.995E+01

0.15 1.9

39.24 41.17

0.28 1.4

2.333 — 2.39E+02 22.21 3.05E02 9.81E+05

— 3.34E+04

HD2294Pb1-Cc

613.9

”

Calcite

5.0

33.30

3.06E02 1.77E02 7.55E+00 1.47E01 6.413E+01

0.16

1.560E+01

0.20

39.61

0.23

4.110 2.47E+02 1.77E+00

4.25E+04

a

Measured from the North portal.

b

Ratios corrected for procedural blank and mass-discrimination. Blank values: [U] = 1 pg (±80%). c Not determined. d

1.183 4.95+02

Sample HD2092 was collected in Alcove 3 about 21.5 m from the main tunnel.

206

Pb/204Pb = 18.76 ± 0.21,

207

Pb/204Pb = 15.58 ± 0.094,

208

Pb/204Pb = 38.02 ± 0.27. Rho (206Pb/204Pb  207Pb/204Pb) = 0.8, [Pb] = 1.5 pg (±80%),

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

HD2155Pb1-Cc HD2177Pb1-Cc HD2177Pb2-Cc HD2177Pb1-Mn HD2177-Mn2

2073

2074

Table 2 U-series and Sr isotope data and Subsample name

234

230

Th/U ages in samples of secondary minerals from the Exploratory Studies Facility (ESF), unsaturated zone, Yucca Mountain, Nevada

238

U/

U ARa

230

Th/238U ARa

234 U/238U initial ARb

n.a.h 166.3 ± 1.3 n.a. n.a. n.a. 361.3 ± 9.8 n.a. n.a. 135.4 ± 0.9 n.a. 497 ± 66 n.a. 366 ± 34 n.a. n.a. n.a. 276 ± 21 94.6 ± 3.7 >500 >500 n.a. n.a. >500

n.a. 5.013 ± 0.012 n.a. n.a. n.a. 7.02 ± 0.12 n.a. n.a. 6.493 ± 0.014 n.a. 5.47 ± 0.81 n.a. 2.39 ± 0.13 n.a. n.a. n.a. 1.448 ± 0.023 4.351 ± 0.070 n.a. n.a. n.a. n.a. n.a.

Mass fraction of adsorbed U(u)c

U sorption distribution coefficient k d (mL g1)d

87

Rb/86Sre

87 Sr/86Sr measuredf

87 Sr/86Sr initialg

0.014–0.041 23–66 n.m.i n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. 0–0.027 0–191 0.061 0.71237 0.71236 0.010–0.036 196–727 n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. 0.218–0.239 450-492 0.064 0.71156 0.71155 0.228–0.248 781–852 n.m. n.m. n.a. 0.032–0.058 44–79 0.066 0.71202 0.71201 0.110–0.134 204–247 n.m. n.m. n.a. 0.049–0.074 19–30 0.94 0.71236 0.71219 0.023–0.049 70–148 n.m. n.m. n.a. 0.029–0.054 71–135 2.1 0.71280 0.71242 0.053–0.078 183–270 n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. 0.044–0.069 72–114 0.94 0.71165 0.71148 0.041–0.067 31–50 n.m. n.m. n.a. n.a. n.a. n.m. n.m. n.a. a Activity ratios (AR) calculated from measured isotope ratios (activity N1/N2 = atomic N1/N2*k1/k2) using decay constants: k238 = 1.55125  1010 y1; k234 = 2.835  106 y1, and k230 = 9.1952  106 y1 for 238U, 234U, and 230Th, respectively. Errors are 2r. b Ages in thousands of years before present (ka) and initial 234U/238U AR are calculated without a correction for detrital U and Th contribution and are ambiguous for the Mn-oxide subsamples. Errors are 2r. c u, A range of values for mass fraction of adsorbed 238U calculated for 234U/238U activity ratios of 0.939 and 0.800 in the solid phase and 6 in the adsorbed phase. d Kd, Range of distribution coefficient for the adsorbed U calculated for minimum and maximum u values. Kd = Cadsorbed/Cwater, where C is concentration. e Values of 87Rb/86Sr are calculated from Rb and Sr concentrations in Table EA1. f Total uncertainty for 87Sr/86Sr values including external error is ±0.00005. g Measured ratio corrected for the in situ 87Rb decay over 12.8-my age of the host tuffs. h n.a., not applicable (age, 234U/238U initial AR, or 87Sr/86Sr initial are not calculable; u and Kd were not calculated for opal and calcite subsamples). i n.m., not measured. Mass fractions of adsorbed U (u) and distribution coefficients Kd are discussed in the Electronic Annex EA2.

HD1939Mn2 HD2055Pb11-Op HD2055Pb12-Op HD2057Pb1-Cc HD2057Pb2-Cc HD2057Pb2-Op HD2062Pb2-Mn HD2062Mn2 HD2074Pb2-Op HD2089APb3-Mn HD2089AMn2 HD2092Pb1-Mn HD2092-Mn2 HD2177Pb1-Mn HD2177-Mn2 HD2227Pb1-Mn HD2227-Mn2 HD2231Pb1-Cc HD2233Pb1-Ch2 HD2233Pb2-Ch2 HD2247Pb1-Mn HD2247-Mn2 HD2293Pb1-Op

1.011 ± 0.005 3.505 ± 0.008 1.003 ± 0.010 3.44 ± 0.33 2.055 ± 0.061 3.781 ± 0.008 0.939 ± 0.006 0.988 ± 0.003 4.740 ± 0.013 2.043 ± 0.004 2.091 ± 0.006 1.100 ± 0.031 1.494 ± 0.004 1.184 ± 0.026 1.057 ± 0.004 1.083 ± 0.008 1.205 ± 0.005 3.563 ± 0.071 1.023 ± 0.016 1.015 ± 0.013 1.161 ± 0.003 1.148 ± 0.010 1.004 ± 0.009

1.123 ± 0.009 3.147 ± 0.010 1.013 ± 0.033 n.m. n.m. 3.168 ± 0.017 n.m. 1.012 ± 0.014 3.835 ± 0.011 n.m. 2.500 ± 0.015 n.m. 1.610 ± 0.012 n.m. 1.241 ± 0.014 n.m. 1.166 ± 0.009 2.256 ± 0.040 1.041 ± 0.018 1.012 ± 0.006 n.m. 1.244 ± 0.010 1.007 ± 0.006

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

230 Th/U age (ka)b

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

2075

in most opal/chalcedony and calcite subsamples are much lower than 238U/204Pb ratios (Table 1), so in situ production of thorogenic 208Pb is small. These results suggest that calcite can be used as a U-series geochronometer, in agreement with previous studies (Szabo and Kyser, 1990; Paces et al., 2001), and probably as a U–Pb geochronometer, but Mnoxides cannot.

corrected for mass fractionation of 0.0010 ± 0.0003 per mass unit using data for Pb-isotope standards SRM-981 and SRM-982 from the National Institute of Standards (NIST) measured at the same run conditions. The mean 234 U/238U in NIST SRM 4321B is 52.879 ± 0.004  106 (n = 134) and 230Th/232Th in the Institute for Reference Materials and Measurements standard IRMM-036 is 31.09 ± 0.04 107 (n = 30). The precision and accuracy of the 234U/238U and 230 Th/238U activity ratios (AR) measured at ROM and USGS were checked with a solution of a secular equilibrium material (70-Ma uranium ore, Ludwig et al., 1985). The AR values were calculated from the measured atomic ratios using known decay constants (Jaffey et al., 1971; Cheng et al., 2000). Results of repeated analyses are reproducible within the 2r error limits, and weighted means of the 234U/238U and 230Th/238U ARs are within analytical error of unity. The Pb and U isotopic data in both laboratories were reduced using the PBDAT program (Ludwig, 1987), which makes corrections for mass-discrimination, procedure blank, spike contribution, and initial Pb.

Mn-oxides have 87Sr/86Sr ratios ranging from 0.71156 to 0.71280 and initial ratios (corrected for in situ 87Rb decay over the 12.8-my age of the host tuff, the maximum possible age of the Mn-oxides) from 0.71148 to 0.71242 (Table 2). Relatively low 87Rb/86Sr ratios in the Mn-oxides ranging from 0.061 to 2.1 (Table 2) result in calculated initial 87 Sr/86Sr that are close to the measured values. Measured 234 U/238U and 230Th/238U ARs in Mn-oxides range from 0.939 ± 0.006 to 2.091 ± 0.006 and from 1.012 ± 0.014 to 2.500 ± 0.015, respectively (Table 2), and are positively correlated.

4. RESULTS

4.3. Pb isotopes in opal/chalcedony, Mn-oxides, and calcite

4.1. Compositions of secondary minerals Compositions of hydrogenic cavity-coating minerals from the Yucca Mountain UZ are highly variable (Table EA1), in contrast with uniform compositions of host tuffs (Peterman and Cloke, 2002). Calcite, the most abundant mineral in cavity coatings in the Yucca Mountain UZ, is depleted in most trace elements, including U, Th, and REE, relative to the host tuffs (Table EA1). Subsamples composed of mixed calcite and opal/chalcedony (Table EA1) have concentrations of most elements similar to lower values measured in calcite subsamples with the exception of U, which is highly enriched. Mn-oxides commonly are depleted in alkaline elements (except Li) and enriched in Ca, Sr, Ba, Mo, Th, U, and Pb relative to host tuff (Table EA1). Calcite subsamples have lower REE concentrations and less-fractionated light REE pattern than the host rocks (smaller chondrite-normalized LaN/SmN, Table EA1), show deep negative Ce anomalies, and have LaN/CeN > 1 (Table EA1). Mn-oxide subsamples are REE-enriched relative to the host tuff, have pronounced positive Ce anomalies, and, except for one, have LaN/CeN < 1. These REE data agree well with previous REE data for Yucca Mountain UZ minerals (Vaniman and Chipera, 1996; Denniston et al., 1997). Concentrations of U, Th, and Pb in the secondary minerals vary by eight orders of magnitude (Tables EA1 and 1, Fig. 4). Calcite is depleted in all three elements relative to the host tuffs, whereas opal/chalcedony and Mn-oxides commonly are enriched in U. Concentrations of Th and initial Pb (204Pb) are high in Mn-oxide and low in most opal/ chalcedony and calcite subsamples (Fig. 4). These data confirm previous observations that Yucca Mountain UZ opal/ chalcedony is suitable for U-series and U–Pb dating (Neymark et al., 2000, 2002; Paces et al., 2001, 2004; Wilson et al., 2003; Nemchin et al., 2006). The 232Th/204Pb ratios

4.2. Sr, U, and Th isotopes in Mn-oxides

Lead isotopic compositions in secondary minerals are highly variable (Table 1). The highest concentrations of radiogenic 206Pb and 207Pb (but not 208Pb) are found in opal/chalcedony, in agreement with previous results (Neymark et al., 2000, 2002). Large 206Pb/204Pb and 207Pb/204Pb in opal/chalcedony are caused by in situ production of radiogenic Pb from U decay in this U-enriched and Th- and initial-Pb-depleted mineral. Most of the opal/chalcedony data points plot below the 12.8-Ma (maximum age) U–Pb isochrons (Figs. 5 and 6) suggesting that the excesses of 206Pb and 207Pb are small or non-existent. Some opal/ chalcedony data points plot above the 12.8-Ma Th–Pb isochron (Fig. 7) showing excess 208Pb. U–Pb and Th–Pb isochrons (loci of coeval samples with the same initial Pbisotope compositions and variable 238U/204Pb, 235U/204Pb, and 232Th/204Pb ratios) are curved in Figs. 5–7 because logarithmic scales are used to show large ranges of isotopic ratios. 238U–206Pb isochrons (Fig. 5) shown as solid lines are calculated assuming initial 234U/238U AR (Ui) of 10 (maximum value of the Ui calculated for opals dated by 230 Th/U method; Neymark and Paces, 2000; Paces et al., 2001). Dotted line isochrons calculated assuming Ui = 1.0 are given in Fig. 5 for ages of 12.8 and 0.01 Ma to emphasize that the effect of initial 234U enrichment on the 206 Pb/204Pb ratio is much larger in younger minerals. In contrast to opal/chalcedony, Pb-enriched Mn-oxides have uniform Pb isotopic compositions (Table 1 and Figs. 5–7) with mean values of 206Pb/204Pb = 18.24 ± 0.18 (n = 6, ±1 standard deviation, SD), 207Pb/204Pb = 15.589 ± 0.013, and 208Pb/204Pb = 39.132 ± 0.050. Median values for these ratios of 18.169 ± 0.011 (±1 median absolute deviation, MAD), 15.592 ± 0.006, and 39.151 ± 0.022, respectively, are very close to the mean values. The 206 Pb/204Pb ratio is slightly larger than, and 207Pb/204Pb and 208Pb/204Pb ratios are very close to the mean Pb isotopic compositions of the host tuffs (206Pb/

2076

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089 1,000

100

1

1

U (μg/g)

1

/U

00

10

U

=

0

1

00

/U

=

0.

Th

0.1

0.

Th

0 .0

/ Th

=

U h/

10

=

T

0.01

0.001 0.001

0.01

0.1

1

10

100

1,000 10,000 100,000 1,000,000

Th (ng/g) 1,000 3

μ=

100

10

7

U (μg/g)

10

10

μ=

1 5

0.1

0

1 μ=

μ=

10

Calcite

0.01

μ= 0.001 0.0001

0.001

0.01

0.1

1

Opal

10

0.1

This work

Mn-Ox Opal [1, 2, 3] Calcite [1]

Average Host Tuff [4]

100

1,000

10,000 100,000

204

Pb (ng/g)

Fig. 4. Plots showing concentrations of Th (A) and 204Pb (B) relative to concentrations of U in calcite, opal/chalcedony, and Mn-oxide subsamples, and average values for host tuff from the Yucca Mountain unsaturated zone. Published data are from [1] Paces et al. (2001); [2], [3], and [4] Neymark et al. (2000, 2002), and 1995. Diagonal lines correspond to Th/U and 238U/204Pb (l) ratios.

204

Pb = 18.123 ± 0.017, ±1 SD; 207Pb/204Pb = 15.578 ± 0.015, and 208Pb/204Pb = 39.15 ± 0.05; n = 15, Neymark et al., 1995). The Pb isotopic composition of Mn-oxides is used in this paper as an initial Pb value in opal/chalcedony and calcite U–Pb age calculations assuming that it reflects the Pb isotopic composition of fracture water at the time of the mineral deposition and that it remains relatively constant throughout the history of the Yucca Mountain UZ. Similar to opal/chalcedony, calcite has highly variable 206 Pb/204Pb from 18.47 ± 0.02 to 23,850 ± 2,000 and 208 Pb/204Pb from 37.53 ± 0.15 to 341 ± 27, but much less variable 207Pb/204Pb from 15.53 ± 0.05 to 16.42 ± 0.03 (Table 1 and Figs. 5–7). Two extreme 207Pb/204Pb values for calcite of 15.35 ± 0.44 and 16.75 ± 0.75 are omitted from this discussion because of their large errors. The mean 207 Pb/204Pb value of 15.67 ± 0.18 (±1 SD, n = 34) is slightly larger than the median value of 15.605 ± 0.024 (±1 MAD). The population of 207Pb/204Pb ratios in these subsamples of calcite consists of a compact main cluster and a ‘‘tail” extending up to 16.42. The four subsamples

of calcite that plot beyond the error above the 12.8-Ma isochron in the 235U/204Pb–207Pb/204Pb plot in Fig. 6B are in this ‘‘tail”. All of these subsamples are closely associated with opal (Table 1 and Neymark et al., 2000, 2002; Paces et al., 2001). The largest 206Pb/204Pb and 208Pb/204Pb ratios are found in several subsamples from outer layers of calcite coating HD2089A (Table 1). The correlation between 206Pb/204Pb and 208Pb/204Pb in this sample is linear with a slope of 0.0126 ± 0.0007. This slope gives an estimate of the 208 Pb/206Pb ratio in the excess Pb component, because contributions from in situ Th- and U-supported 208Pb and 206 Pb are negligible. Measured 206Pb/204Pb and 208Pb/204Pb ratios anti-correlate with 204Pb concentrations in this sample so that the largest 206Pb/204Pb and 208Pb/204Pb values are observed in subsamples with the smallest 204Pb concentrations. The 206Pb/207Pb ratios measured by ICP-MS in seven different calcite subsamples from HD2089A range from 14.3 to 38.8 (Table EA1) and do not correlate with Ba concentrations measured in the same subsample digestions (R2 = 0.005).

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

2077

204

10 6 yr

10 4 yr

1000

206

Pb/

10 6 yr 10 5 yr

Opal (published)

Pb

10000

This work

12 .8

Opal

x

Calcite MnOx

10 9 yr

10 10 yr

100000

100

10 1.E-01 1.E+00 1.E+01 1.E+02 1.E+03 1.E+04 1.E+05 1.E+06 1.E+07 1.E+08 1.E+09 238

204

U/

Pb

Fig. 5. 238U/204Pb–206Pb/204Pb isochron plot for calcite, opal/chalcedony, and Mn-oxides from the Yucca Mountain unsaturated zone. Both axes are logarithmic so theoretical isochrons for closed system evolution (shown for ages between 104 and 1010 years for an initial 234U/238U activity ratio of 10) appear as curves rather than straight lines. For ages 104 and 12.8  106 years dotted lines represent isochrons for an initial 234 U/238U activity ratio of 1. Most opal/chalcedony data plot below and almost all calcite data plot above, the 12.8  106 isochron corresponding to the age of the host tuff. Error bars (±2r) are shown when they exceed the symbol size.

4.4.

230

Th/U and U–Pb ages

4.4.1. 230Th/U ages and initial 234U/238U activity ratios Opal subsamples from outer (younger) portions of coatings HD2055, HD2057, and HD2074 have 230Th/238U ages from 135.4 ± 0.9 to 361.3 ± 9.8 ka and Ui of 5.013 ± 0.012 to 7.02 ± 0.12 (Table 2). Measured 234U/238U and 230 Th/238U ARs are close to unity in opal/chalcedony from older portions of coatings HD2055 (Pb12), HD2233, and HD2293 (Table 2) and indicate ages >500 ka and no recent U mobility. Of the calcite subsamples, only HD2231Pb1-Cc has measurable 230Th and yields a 230Th/238U age of 94.6 ± 3.7 ka and Ui of 4.351 ± 0.070 (Table 2). Reported opal and calcite ages are calculated without correction for detrital U and Th, because very low 232Th (Table 1) renders such correction insignificant. These U-series data are in agreement with published results for opal and calcite from the Yucca Mountain UZ (Neymark and Paces, 2000; Neymark et al., 2000; Paces et al., 2001, 2004). Three Mn-oxide subsamples yield 230Th/238U apparent ages of 276 ± 21 to 497 ± 66 ka and Ui values of 1.448 ± 0.023 to 5.47 ± 0.81 (Table 2). These ages are calculated assuming a closed system without correction for detrital U and Th and are ambiguous because of large 232 Th concentrations (Table 1) and U mobility. Another four Mn-oxide subsamples where both U and Th isotopic compositions were measured contain slight excesses of 230 Th that did not allow age calculations. 4.4.2. U–Pb ages and initial 234U/238U activity ratios of opal/ chalcedony Relatively precise 206Pb/238U and 207Pb/235U ages were calculated for all opal/chalcedony subsamples (Table 3). These ages are reversely discordant (206Pb/238U age > 207 Pb/235U age) indicating the presence of excess 206Pb, in

agreement with previous results (Neymark et al., 2000, 2002; Wilson et al., 2003; Nemchin et al., 2006). The 207 Pb/235U ages range from 0.1329 ± 0.0080 to 9.10 ± 0.21 Ma (Table 3). The Ui values required to make 206 Pb/238U and 207Pb/235U ages concordant (to support observed excess 206Pb) range from 4.22 ± 0.13 to 19.95 ± 0.22 (Table 3) and most of these values are within the range of 2.0–10.0 observed for the 234U/238U ARs in Yucca Mountain UZ pore water (Paces and Neymark, 2004) and calculated Ui values from U-series data for the UZ opal (Neymark et al., 2000; Paces et al., 2001). However, Ui > 10 may indicate either (1) the presence of waters with 234 U/238U AR up to about 20 in the Yucca Mountain UZ in the past or (2) the presence in opal/chalcedony of 206Pb unsupported by in situ decay of both 238U and 234U. The second possibility is more plausible because two of three subsamples where Ui values were calculated from both 230 Th/U and U–Pb data (HD2055Pb11-Op, HD2057Pb2Op, and HD2074Pb2-Op, Tables 2 and 3) have Ui values estimated from U–Pb data that are almost two times larger than Ui values calculated from 230Th/U data (9.38 ± 0.24 versus 5.013 ± 0.012 and 11.0 ± 1.9 versus 6.493 ± 0.014). The largest opal/chalcedony apparent Ui value of 19.95 ± 0.22 (Table 3) is calculated for subsample HD2233Pb1-Ch1 collected from the outer portion of the coating. The U–Pb data for this subsample plot above the Ui = 10 disequilibrium concordia curve (Fig. 8A). This subsample also has the largest 208Pb/204Pb ratio measured in this sample (Table 1), which does not depend on the initial 234 U excess, thus indicating the presence of U- and Thunsupported 206Pb and 208Pb. The 206Pb/238U ages plotted against the distance of the subsampled interval from the surface of this coating correlate with the microstratigraphy (except for two subsamples near the surface) of the coating

2078

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

MnOx Opal (published)

207

Pb/

204

Pb

60

This work

105 yr

Opal

70

106 yr

Calcite

12.8 x 10 6 yr

80

108 yr

90

50 40

30

20

10 1.E-03

1.E-01

1.E+01 235

1.E+03

1.E+05

204

U/

Pb

17.0

16.0

207

Pb/204Pb

16.5

15.5

15.0 1.E-02

1.E-01

1.E+00 235

1.E+01

1.E+02

204

U/

Pb

Fig. 6. (A) 235U/204Pb–207Pb/204Pb isochron plot for calcite, opal/chalcedony, and Mn-oxides from the Yucca Mountain unsaturated zone. The x-axis is logarithmic so theoretical isochrons for closed system evolutions (shown for ages between 105 and 108 years) appear as curves rather than straight lines. Most opal/chalcedony data plot below the 12.8  106 isochron indicating a younger age than the host tuff. Most calcite data and all Mn-oxide data plot on the horizontal line corresponding to the initial-Pb 207Pb/204Pb ratio. (B) Dotted rectangle from (A) shown at a larger scale. Four data points plotting above the 12.8  106 isochron are calcite subsamples closely associated with opal. Error bars (±2r) are shown when they exceed the symbol size.

(Fig. 8B) and define a long-term average growth rate of 1.1 ± 0.2 mm my1 The 207Pb/235U ages, which are systematically younger than the corresponding 206Pb/238U ages, correlate with the microstratigraphy through the entire depth and define the same long-term average growth rate of 1.2 ± 0.3 mm my1. A similar long-term average growth rate of about 2 mm my1 is calculated for the outer half of coating HD2055. Subsamples -Pb11 and -Pb12 in this sample, collected at depths of about 0.5 and 9 mm, respectively,

207 yield Pb/235U ages of 4.702 ± 0.020 Ma (Table 3).

0.2451 ± 0.0034

and

4.4.3. U–Pb ages and initial 234U/238U activity ratios of calcite Most 206Pb/238U ages determined for the calcite subsamples are much older than the 12.8-Ma age of the host tuff (Table 3 and Fig. 5) and thus unreasonable. Apparent 208 Pb/232Th ages of all but two samples also exceed the

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

2079

400

10 10 yr

350

208

204

Pb/ Pb

300 250 200 150

Opal (measured) Opal (calculated) Calcite (measured) Calcite (calculated) MnOx

This work

Opal (published)

100 9

10

50 0 0.0001

0.001

0.01

0.1 232

10

1

yr

100

1000

Th/204Pb

75 70

208

Pb/204Pb

65 60 55 50 45

10

10

yr

40 35 0.0001

0.001

0.01 232

232

204

208

0.1

1

10

204

Th/ Pb

204

Fig. 7. (A) Th/ Pb– Pb/ Pb isochron plot for calcite, opal/chalcedony, and Mn-oxides from the Yucca Mountain unsaturated zone. The x-axis is logarithmic so theoretical isochrons for closed system evolution (shown for ages between 12.8  106 and 1010 years) appear as curves rather than straight lines. (B) Dotted rectangle from (A) shown at a larger scale. Error bars (±2r) are shown when they exceed the symbol size.

age of the host tuff (Table 3 and Fig. 7). In contrast, most 207 Pb/235U ages for calcite are indistinguishable from zero within error (Table 3). Calcite subsamples HD2055Pb11, HD2055Pb12, HD2059Pb4,2 HD2074Pb2 (-Cc1 and -Cc2), and HD2293Pb1, which are closely associated with opal and have elevated 207Pb/204Pb, yield non-zero but unreasonable 207Pb/235U ages from 10.9 ± 5.3 to 162 ± 25 Ma that are substantially older than 207Pb/235U ages of coexisting opal from 0.1329 ± 0.0080 to 4.702 ± 0.020 Ma (Table 3). Apparent Ui values from 1388 ± 6 to 3315 ± 86 (Table 3) are calculated from measured206Pb/238U ratios and 207Pb/235U ages of associated opal for five calcite subsamples assuming closed U–Pb system behavior. These unreasonably large Ui values are required to support observed excess 206Pb if it is caused by in situ decay of 2 Opal from sample HD2059 was analyzed previously (Neymark et al., 2000, 2002; Paces et al., 2001).

initial excess 234U; they indicate that initial excess of in calcite cannot explain observed excesses of 206Pb.

234

U

4.4.4. Isotopic composition of initial Pb in calcite Isotopic composition of initial Pb in calcite of unknown age can be evaluated using two approaches: (1) assuming that it is the same as in coexisting low-238U/204Pb minerals (Mn-oxides) and (2) correcting for the radiogenic Pb produced in situ from U and Th decay assuming closed U–Th–Pb system behavior in calcite and using age and initial 234U/238U AR in dated coeval opal. The first approach produces either unrealistically old U–Pb ages of calcite or extremely large Ui values (see previous section). The second approach is applied to five calcite subsamples where coexisting opal also has been analyzed (Table 4). Apparent initial 206 Pb/204Pb and 207Pb/204Pb are slightly lower than the measured values and vary from 36.7 ± 3.4 (2r) to 316 ± 1.5 and from 15.573 ± 0.031 to 16.346 ± 0.034, respectively. These apparent initial 206Pb/204Pb and

2080

Table 3 Radiogenic Pb/U and Pb/Th ratios, U–Pb and Th–Pb ages, and calculated initial unsaturated zone, Yucca Mountain, Nevada Pb*/238Ua

Pb*/235Ua

234

U/238U activity ratios in samples of secondary minerals from the Exploratory Studies Facility (ESF),

Mineral

206

Error (%)

207

Error (%)

208

Pb*/232Tha

Error (%)

206

Pb*/238U Age (Ma)b

207

Pb*/235U Age (Ma)b

208

Pb*/232Th Age (Ma)b

234

U/238 Uini ARc

HD1939Pb1-Cc HD2055Pb6-Cc HD2055Pb7-Cc1 HD2055Pb7-Cc2 HD2055Pb10-Cc HD2055Pb11-Cc HD2055Pb11-Op HD2055Pb12-Cc HD2055Pb12-Op HD-2057-Pb1-Cc HD-2057-Pb2-Cc HD-2057-Pb2-Op HD2059Pb4-Cc HD2062Pb1-Cc HD2062Pb2-Cc HD2065Pb4-Cc HD2074Pb1-Cc2 HD2074Pb1-Cc3 HD2074Pb2-Cc1 HD2074Pb2-Cc2 HD2074Pb2-Op HD2089APb1-Cc1 HD2089APb1-Cc2a HD2089APb1-Cc2b HD2089APb1-Cc2c HD2089APb1-Cc3 HD2089APb2-Cc HD2092Pb1-Cc HD2098Pb3-Cc HD2109Pb1-Cc HD2155Pb1-Cc HD2177Pb1-Cc HD2177Pb2-Cc HD2227Pb1-Cc HD2231Pb1-Cc HD2233Pb1-Ch1e HD2233Pb1-Ch2 HD2233Pb2-Ch1 HD2233Pb2-Ch2

Calcite Calcite Calcite Calcite Calcite Calcite Opal Calcite Opal Calcite Calcite Opal Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Opal Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Chalcedony Chalcedony Chalcedony Chalcedony

5.951E03 1.317E01 1.029E01 7.102E02 1.624E02 5.830E02 1.700E04 1.279E01 8.886E04 2.532E01 3.935E02 2.255E04 3.094E01 6.736E02 1.860E02 8.177E02 1.224E01 1.270E02 2.286E02 1.221E02 8.264E05 2.427E+00 1.687E+00 2.203E+00 1.654E+00 9.872E01 6.618E01 5.834E03 5.289E01 6.822E01 6.378E02 2.199E+01 1.039E02 4.438E02 4.818E04 1.674E03 1.196E03 1.083E03 1.063E03

9.8 15 3.8 4.9 2.8 0.41 0.43 0.18 0.41 1.7 0.51 0.24 0.33 0.18 3.2 0.23 0.27 0.59 0.24 0.41 0.59 0.62 0.20 0.95 0.45 0.37 0.32 2.3 0.62 0.77 0.35 0.35 1.1 0.34 17 0.37 0.30 0.70 0.30

2.774E02 — 6.497E02 3.243E01 2.303E02 2.927E02 2.415E04 4.944E02 4.641E+00 — — 3.885E04 1.735E01 3.242E03 — 1.852E02 — 2.860E03 1.183E02 1.075E02 1.309E04 1.624E02 — 1.487E02 9.163E+00 — — 9.804E03 — 6.132E01 3.013E04 3.013E01 6.352E03 — 1.223E02 4.162E03 4.098E03 4.948E03 4.959E03

223 — 624 118 168 6.0 1.4 3.9 0.43 — — 0.59 15.3 25 — 46 — 200 17 48 6.1 72 — 60 112 — — 150 — 79 5770 166 214 — 81 2.1 1.08 2.5 1

4.781E03 — — 9.663E02 — — 1.013E01 — 4.348E03 — — 4.042E02 8.967E01 — 4.486E04 1.376E02 — 6.655E02 — — — — — — — 3.108E+00 8.272E03 1.820E03 1.248E02 7.977E03 3.044E02 7.995E02 2.425E03 5.541E03 1.602E04 — 5.689E+01 — 1.654E+01

70 — — 85 — — 32 — 18 — — 41 31 — 674 192 — 289 — — — — — — — 27 18 17 44 26 20 9.0 488 22 234 — 220 — 70

38.2 ± 3.7 798 ± 119 631 ± 24 442 ± 22 103.8 ± 2.9 365.3 ± 1.5 1.096 ± 0.005 775.9 ± 1.4 5.726 ± 0.024 1455 ± 25 248.8 ± 1.3 1.454 ± 0.004 1737.9 ± 5.7 420.2 ± 0.8 118.8 ± 3.8 505.9 ± 1.6 744.6 ± 2.0 81.37 ± 0.48 145.70 ± 0.35 78.26 ± 0.32 0.5327 ± 0.0032 7940 ± 49 6372 ± 13 7504 ± 71 6292 ± 29 4423 ± 30 3270 ± 14 37.50 ± 0.85 2734 ± 20 3353 ± 26 398.6 ± 1.4 20209 ± 50 66.62 ± 0.73 279.9 ± 1.0 3.11 ± 0.53 10.781 ± 0.039 7.687 ± 0.032 6.979 ± 0.049 6.85 ± 0.021

28 ± 62 — 64 ± 399 285 ± 336 23 ± 39 29.3 ± 1.8 0.2451 ± 0.0034 49.0 ± 1.9 4.702 ± 0.020 — — 0.3944 ± .0023 162 ± 25 3.29 ± 0.83 — 18.8 ± 8.7 — 2.9 ± 5.8 11.9 ± 2.0 10.9 ± 5.3 0.1329 ± 0.0080 16 ± 12 — 15 ± 9 9 ± 10 n.d. n.d. 10 ± 15 — 486 ± 383 0 ± 18 277 ± 460 6 ± 14 — 12 ± 10 4.218 ± 0.08 4.133 ± 0.093 5.01 ± 0.12 5.023 ± 0.034

96 ± 68 — — 1864 ± 1589 — — 1951 ± 630 — 88 ± 15 — — — 12900 ± 4040 — 9 ± 61 276 ± 537 — 1300 ± 3800 — — — — — — — 28600 ± 7700 166 ± 30 37 ± 6 250 ± 110 161 ± 42 607 ± 123 1555 ± 140 49 ± 239 112 ± 24 3.2 ± 7.5 — 82030 ± 180500 — 57900 ± 40800

—d — — — — 3315 ± 86 9.38 ± 0.24 2326 ± 8 4.22 ± 0.13 — 1340 ± 1 7.17 ± 0.04 — — — — — — 1645 ± 298 3078 ± 568 11.0 ± 1.9 — — — — — — — — — — — — — 19.95 ± 0.22 11.46 ± 0.13 6.90 ± 0.28 6.49 ± 0.12

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

Subsample name

Pb/204Pb = 39.132 ± 0.052.

HD2233 ESF 6838.1m

9 Ma 8 Ma

-Pb1-Ch1

7 Ma

0.0016

6 Ma

0.0014

5 Ma

206

Pb*/ 238U

0.0018

208

0.0012

4 Ma

0.0010

e

c

d

10 8 6 4 2 1 U i= U i= U i = U i = U i = U i =

0.0008 0.002

0.004

0.006 207

Pb*/

0.008

0.010

235

U

0

Distance from Coating Surface (mm)

Values corrected for common Pb (*) using weighted average of Mn-oxide values (Table 2): 206Pb/204Pb = 18.163 ± 0.024, 207Pb/204Pb = 15.589 ± 0.013, and Conventional U–Pb ages calculated assuming initial radioactive equilibrium. Initial activity ratio (AR) calculated from 207Pb*/235U age and measured 206Pb*/238U assuming closed system behavior. Not determined. HD2233 subsamples -Pb1 through -Pb5 were collected at distances 0 to 0.3, 1.0 to 2.0, 2.0 to 4.0, 4.0 to 6.0, and 6.0 to 8.0 mm from the coating surface. a

2081 10 Ma

0.0020

b

HD2233Pb3-Ch HD2233Pb4-Ch HD2233Pb5-Ch HD2247Pb1-Cc1 HD2247Pb1-Cc2 HD2247Pb2-Cc HD2293Pb1-Cc HD2293Pb1-Op HD2294Pb1-Cc

Chalcedony Chalcedony Chalcedony Calcite Calcite Calcite Calcite Opal Calcite

1.237E03 1.597E03 1.816E03 2.332E02 6.455E03 2.243E02 7.638E02 8.767E04 2.568E01

0.99 0.65 2.4 3.4 4.1 1.6 0.23 0.50 0.27

5.706E03 9.004E03 8.852E03 1.563E02 — 5.443E03 8.307E02 3.386E03 1.140E02

3.4 2.4 8.4 548 — 799 17 0.94 213

— — — 8.794E03 3.431E03 3.263E02 — 5.713E02 1.073E02

— — — 105 77 59 — 28 20

7.972 ± 0.079 10.284 ± 0.067 11.69 ± 0.27 148.6 ± 5.0 41.5 ± 1.7 143.0 ± 2.3 474.5 ± 1.1 5.649 ± 0.028 1473.4 ± 4.0

5.78 ± 0.19 9.10 ± 0.21 8.95 ± 0.75 16 ± 86 — 6 ± 44 81 ± 13 3.433 ± 0.032 12 ± 25

— — — 177 ± 186 69 ± 53 649 ± 386 — 1120 ± 320 216 ± 44

7.54 ± 0.45 4.68 ± 0.38 9.1 ± 1.6 — — — 1388 ± 6 7.60 ± 0.16 —

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

1

Data are not included in growth rate calculations

2 3

y = (1.2±0.3)x - 4.3 2 R = 0.92

4

y = (1.1±0.2)x - 6.1 2 R = 0.99

5 207

6 206

235

Pb / U 238 Pb / U

7 8 0

5

10

15

Age (Ma) Fig. 8. U–Pb data for chalcedony subsamples from sample HD2233. (A) Disequilibrium concordia diagram of common-Pbcorrected 206Pb*/238U and 207Pb*/235U ratios shown with error ellipses. Multiple concordia curves correspond to different initial 234 U/238U activity ratios (Ui). (B) 207Pb/235U and 206Pb/238U ages for subsamples from HD2233 plotted against distance from the coating surface (microstratigraphic depth).

207

Pb/204Pb are larger than values measured in low-238U/ Pb Mn-oxides and indicate the presence of excess 206Pb and 207Pb unsupported by 238U, 234U, and 235U decay in these calcite subsamples. The 206Pb/207Pb ratios of 134 ± 43 (2r) to 391 ± 24 (Table 4) in this calcite that has U-unsupported Pb are calculated by correcting apparent initial ratios with ratios measured in Mn-oxides. 204

5. DISCUSSION 5.1. Mobility of Sr and U isotopes 5.1.1. Sr isotopes Strontium isotope ratios in the UZ hydrogenic minerals are inherited from the percolating water at the time of mineral deposition and should remain almost unchanged in Sr-rich calcite and Mn-oxides. Previous studies showed distinct correlations of Sr isotope ratios in calcite with

2082

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

Table 4 Apparent initial 206Pb/204Pb, 207Pb/204Pb, and corrected 206Pb/207Pb in samples of calcite associated with opal from the Exploratory Studies Facility (ESF), unsaturated zone, Yucca Mountain, Nevada 206

Pb/204Pb initiala

Subsample name

Apparent

HD2055Pb11-Cc HD2055Pb12-Cc HD2057Pb2-Cc HD2074Pb2-Cc1 HD2074Pb2-Cc2 HD2293Pb1-Cc

183.62 ± 0.97 316 ± 1.5 41.35 ± 0.08 40.00 ± 0.15 88.82 ± 0.49 36.7 ± 3.4

Apparent

207

Pb/204Pb initialb

16.184 ± 0.037 16.346 ± 0.034 15.573 ± 0.031 15.723 ± 0.067 15.848 ± 0.045 15.726 ± 0.023

Excess

206

Pb/207Pbc

277 ± 23 391 ± 24 n.a.d 161 ± 96 272 ± 61 134 ± 43

a Measured isotope ratios from Table 2 minus in situ radiogenic 206Pb calculated using 207Pb/235U age and initial 234U/238U AR in associated opal from this table. b Measured isotope ratios from Table 2 minus in situ radiogenic 207Pb calculated using 207Pb/235U age of associated opal from this table. c Ratio of apparent initial 206Pb/204Pb and 207Pb/204Pb corrected for common Pb using weighted average of Mn-oxide values (Table 2). [(206Pb/207Pb)corr = [(206Pb/204Pb)ap ini  (206Pb/204Pb)MnOx]/[(207Pb/204Pb)ap ini  (207Pb/204Pb)MnOx]. d Not applicable because measured 207Pb/204Pb is within error of the weighted average 207Pb/204Pb value in Mn-oxides (Table 1).

relative positions of subsamples in a mineral coating (Marshall and Futa, 2001; Paces et al., 2001). Basal (old) calcite generally has less radiogenic 87Sr/86Sr ratios between 0.7105 and 0.7120 than late calcite (0.7115 and 0.7127) and pore water (0.7122 to 0.7127) (Marshall and Futa, 2001; Marshall and Futa, 2003). Johnson and DePaolo (1994) explained systematic variations in calcite 87Sr/86Sr ratios by radiogenic 87Sr ingrowth in the porous and reactive PTn tuffs overlying the TSw rocks that host studied samples. The Mn-oxides, which typically underlie the oldest calcite, show the same range of 87Sr/86Sr ratios as the youngest late-stage calcite. This indicates that the Mn-oxides either formed recently (which contradicts their microstratigraphic position), or more likely behave as open systems and may have been exchanging Sr with, or scavenging it from, percolating water over the history of Yucca Mountain. Therefore, it can be expected that Mn-oxides would retard the migration of 90Sr released from the proposed repository and transported by water percolating through the UZ. 5.1.2. U isotopes Similar to Sr isotope ratios, 234U/238U ARs in secondary hydrogenic minerals in the UZ are inherited from the percolating water. Because the UZ water is 234U-enriched, 234 U/238U AR are >1 at the time of deposition, but approach unity (secular equilibrium) after several half-lives of 234U if the minerals remain closed to mass transfer. However, water–mineral interaction along fracture flow paths causes differential U isotope mobility and radioactive disequilibrium in both the water and the mineral. The degree of disequilibrium depends on water velocity, mineral dissolution rates, a-recoil processes, adsorption and desorption, redox conditions, and precipitation of newly formed minerals (Porcelli and Swarzenski, 2003, p. 321 and references therein). Observed U-series isotope disequilibrium in Mnoxides (Table 2) indicates that these minerals were open to U migration, as well as Sr migration, over the past several hundred thousand years. These data indicate that Mn-oxides exchange U and Sr with percolating water and cannot be used as a reliable dating tool. 234U/238U ARs in Mn-oxides can be used to estimate time-integrated

in situ distribution coefficients Kd for U sorption ranging from 0 to 852 mL g1 (Table 2) as described in the Electronic Annex 2 (EA2). 5.2. Mobility of U and Th decay products: evidence from Pb isotopes Radioactive decay of 238U, 235U, and 232Th contained in the rock mass and cavity-coating minerals produces intermediate decay products, a portion of which can be released from the solids to the cavity via diffusion and a-recoil. The diffusion-related release is mainly important for gaseous Rn isotopes and is negligible for other intermediate products in the decay chains. 226Ra, 224Ra, and 223Ra in solids produce 222Rn, 220Rn, and 219Rn, respectively. The fraction of Rn released from solids by diffusion is a function of the temperature-dependent diffusion coefficient of Rn (Anderson et al., 1963; Coombs and Cuddihy, 1983). At a constant temperature the mean diffusion range RD (cm) of Rn is given by (Anderson et al., 1963; Voltaggio et al., 2001): RD ¼ ðDT  sÞ1=2 ;

ð1Þ 2 1

where DT is Rn diffusion coefficient (cm s ) at temperature T, and s is a mean life [(half-life)/ln 2] of a Rn isotope. Using Eq. (1) and the DT value in air at room temperature of 0.1 cm2 s1 (Tanner, 1964), RD values in air are 2.2 m, 2.8 cm, and 0.76 cm for 222Rn, 220Rn and 219Rn, respectively. Therefore, short-lived Rn isotopes 220Rn and 219Rn released by diffusion migrate over very short distances and may only affect Pb isotope compositions of secondary minerals grown in small cavities. The process of a-decay in U and Th decay chains causes ejection of daughter nuclides from their parent positions. The energies of the a-recoiled nuclides of 70–160 keV (Hashimoto et al., 1985) are sufficient for their displacements (recoil range) in solids up to 102 nm (Jonckheere and Gogen, 2001; Stu¨bner and Jonckheere, 2006). If a-emitters are within the recoil range of the surface of U- and Thbearing minerals, daughter nuclides are ejected directly into the surroundings and may be implanted into adjacent cavity-coating minerals.

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

5.2.1. Excess Pb caused by mobility of intermediate decay products of U and Th Excess radiogenic Pb, unsupported by in situ U and Th decay may originate from decay of intermediate radioactive daughter isotopes incorporated into a mineral during and after its formation. UZ water is enriched in 234U due to a-recoil-related processes and some excess 206Pb is produced by decay of this initial 234U excess. However, 234 U-related excess 206Pb is insufficient to explain the observed large 206Pb/204Pb and 206Pb/207Pb values in calcite in the Yucca Mountain UZ (comparison of isochrons for different Ui is shown in Fig. 5). The most striking feature of the Pb isotopic compositions in calcite is a contrast between highly elevated 206Pb/204Pb and 208Pb/204Pb ratios and a fairly constant and low 207Pb/204Pb ratio. The largest Pb isotope anomaly is in calcite subsample HD2089APb1Cc2b, where measured 206Pb/207Pb reaches a value of about 1400 (Table 1), which is substantially larger than a value of 21.7 for this ratio in radiogenic Pb produced by decay of modern U. To our knowledge, 206Pb/207Pb ratios that high have never been reported for natural materials. These data indicate an extreme decoupling of 238U and 232Th decay chains from the 235U decay chain. The observed correlation between excess 208Pb and 206Pb indicates a common cause for excesses of both isotopes, perhaps excesses of isotopes of Ra and Rn, which occur in both 238U and 232Th decay chains. Calcite may contain initial Ra, a close geochemical analogue for Ba. Both elements should behave similarly during distribution-coefficient-driven coprecipitation with Ca (Gnanapragasam and Lewis, 1995; Curti, 1999). However, the absence of correlation between measured 206Pb/207Pb ratios and Ba concentrations in calcite from coating HD2089A (subsamples Ba1–Ba7; Table EA1) suggests that Ba and the excess Pb in the UZ calcite probably originate from different sources. Median concentrations of Ba and excess 206Pb in these calcite subsamples are about 0.7 lg g1and 24 ng g1, respectively. If excess 206Pb is produced by the decay of excess 226Ra co-precipitated with Ba, then the initial Ba /226Ra ratio in this calcite should be about 30. This estimate is many orders of magnitude smaller than Ba/226Ra values of about 1011 reported for young Ra-enriched hydrothermal calcites (Rihs et al., 2000) where 226 Ra/238U AR is elevated relative to travertine calcite (Eikenberg et al., 2001a). Therefore, initial 226Ra excesses in the UZ calcite probably are not the main cause for the large excess of 206Pb. Excess 206Pb in secondary minerals also can derive from decay of 210Pb produced by a-decay of shorter-lived daughters of 222Rn. For example, young calcite from several caves showed 210Pb/226Ra AR >1 (Baskaran and Iliffe, 1993; Tanahara et al., 1998; Paulsen et al., 2003; Woo et al., 2005). In contrast to caves, excess 210Pb was not found in modern calcite formed in well-aerated environments (Krishnamurthy et al., 2003) or in hydrothermal calcites (Rihs et al., 2000; Condomines and Rihs, 2006). The discussion below shows that Rn-produced 210Pb may be sufficient to explain the excess 206Pb observed in the Yucca Mountain UZ calcite. The presence of excess Pb isotopes derived from Rn decay products is easier to identify in calcite than in U-rich

2083

opal/chalcedony, because concentrations of U- and Th-supported in situ 206Pb, 207Pb, and 208Pb are minimal in calcite. However, obvious excesses of 206Pb and 208Pb were found in the outer portion of chalcedony coating HD2233 (Table 1 and Fig. 8). The 207Pb/235U age of the outer portion of this coating is about 4 Ma (Table 3) thus allowing implantation of Rn decay products into the surface of the coating for a longer time compared to coatings with younger outer portions. 5.2.2. Excess 207Pb and 206Pb in calcite associated with opal Calcite and opal are intimately intergrown in the UZ coatings where both minerals are present (Fig. 2) and have U concentrations differing by orders of magnitude. Therefore, direct a-recoil of a small portion of U daughters from U-enriched opal may cause appreciable 206Pb and 207Pb excesses in associated calcite and only minor complementary deficits in opal. About 2.2% of 234U produced by decay of parent 238U is lost from a 1-lm-radius spherical opal particle via a-recoil of 234Th (Eq. (1) in EA2) and a part of it may be implanted into the host calcite creating elevated 234U/238U AR in the calcite. The 234U/238U AR as high as 410 has been reported for U-poor phases associated with the uranium mineral carnotite and was explained by the a-recoil (Sheng and Kuroda, 1986). Similar proportions of 230Th and 231Pa, which are decay products of 234U and 235 U, respectively, also will be a-recoiled into calcite. Concentrations of these a-recoiled isotopes are proportional to activities of their parents in opal. Decay of these implanted isotopes produces excess206Pb and 207Pb unsupported by in situ decay of 238U and 235U. Calculated 206Pb/207Pb ratios in the excess Pb (Table 4) are substantially larger than 238 U/235U AR of 21.7 in modern U. Calcite closely associated with opal also has elevated 207Pb/204Pb ratios and 207 Pb/235U ages (Table 3 and Fig. 6). These data indicate that both 206Pb/238U and 207Pb/235U ages of U-poor minerals closely associated with fine grains of U-enriched minerals may be grossly overestimated and should be interpreted with caution. 5.2.3. Comparison of radon-produced excess 206Pb in calcite from Yucca Mountain UZ and from speleothems We have considered several possible sources of Pb isotopes unsupported by in situ U and Th decay. Excess of Ra is ruled out as a major source of unsupported Pb in calcite; direct recoil implantation may be important for calcite closely intergrown with micron-size particles of U-rich opal but not for opal-free calcite; excess of Rn and its daughter nuclides may be important in many cases. Substantial concentrations of radon are common in air in the isolated voids in rocks. 222Rn concentrations in cave air vary from hundreds to thousands of Bq m3 (Papachristodoulou et al., 2004 and references therein). Elevated concentrations of 222 Rn (up to 65,000 Bq m3) were measured in the ESF air during periods without ventilation (Unger et al., 2004). The excess 210Pb activity in a mineral A210 Pbex (in disintegrations per minute per gram, dpm g1) produced from 222 Rn decay in cavity water and air is equal to the total 210 Pb activity minus in situ 226Ra-supported activity. In this model the initial 210Pb activity unsupported by 226Ra decay

L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

2

1

ð2Þ 210

where Q(t) is the flux (dpm cm y ) of excess Pb to the calcite–water interface, w(t) is the calcite deposition rate (g cm2 y1) and k210 Pb ¼ 0:0311 y1 is the 210Pb decay constant (Appleby and Oldfield, 1992, p. 731). Formula (2) is equivalent to the equation given by Smith and Ellis (1982) for the 210Pb excess in fluvial sediments. The concentration of excess 206Pbex (atoms g1) produced from this excess 210Pb after conversion of dpm to dpy (disintegrations per year) is:

Pb (ng/g)

The concentration of 206Pbex reaches a steady-state value within 100 y of the mineral formation (Fig. 9A) and in the case of constant 210Pb flux is determined only by the deposition rate so that larger concentrations will build up in a slower growing mineral. Calculated 206Pbex for published 210Pb data in speleothem calcite (Baskaran and Iliffe, 1993; Tanahara et al., 1998; Paulsen et al., 2003) are shown in Fig. 9B together with the data for the Yucca Mountain UZ calcites. Growth rates for these speleothems range from 0.03 to 5.9 mm y1 and are typical for cave deposits elsewhere (Hill and Forti, 1997). The growth rates were recalculated to g cm2 y1 using calcite specific gravity of 2.71 g cm3 to obtain estimates of 206Pbex from formula (3). Estimated 206Pbex values in the speleothem calcites range from about 106 to 103 ng g1. The largest of these concentrations is 200 times smaller than the concentration of radiogenic 206Pb reported by Richards et al. (1998) for a Quaternary speleothem calcite. In contrast to speleothems, excesses of 206Pb are large in slowly grown UZ calcite (Fig. 9B). The growth rate for UZ calcite is estimated from the total coating thickness (Table 1) assuming continuous growth of coatings at a constant rate throughout the 12.8 Ma history of the Yucca Mountain UZ. Most estimated growth rates are very slow (between 107 and 106 mm y1) and are in good agreement with the growth rates of 1  107 to 5  106 mm y1 determined for the UZ coatings using ages of opal/chalcedony intercalated with calcite (Neymark et al., 2000; Neymark et al., 2002; Paces et al., 2004; Nemchin et al., 2006). Concentrations of excess 206 Pb in the UZ calcite indicate that this mineral can scavenge 100 to 102 ng g1 of 210Pb, fluxes of which in the UZ cavities may range from 0.001 to 1 dpm cm2 y1. These fluxes are similar to or slightly smaller than those observed in caves (Fig. 9B) indicating that the 222Rn emanation alone can explain the observed excesses of 206Pb in slowly grown Yucca Mountain UZ calcite. It is also clear that the Rn emanation should not cause an appreciable increase in 206Pb–238U ages of faster growing U-enriched speleothem calcites, but may be a concern for dating slower growing U-poor calcites.

10-1

mm 10 -5

10 10

-3

-1

×y

× mm

-1

× mm

10

-1

y

-3

10

-5

×y

-1

×y mm 0 1 e Rat Q = 1 dpm × cm-2 × y-1 wth -1

y

1

Gro -2

10

-1

10

0

10

1

10

2

10

3

10

4

10

5

Time (y)

206

ð3Þ

mm 10 -7

10

Pbex ¼ k1 210 Pb  525949  ½1  exp ðk210 Pb  t Þ  QðtÞ=wðtÞ:

-1

101

10-7

Excess

206

Pb (ng/g)

A210 Pbex ðtÞ ¼ exp ðk210 Pb  tÞ  QðtÞ=wðtÞ;

103

206

is proportional to the 210Pb supply rate and varies inversely with the mass accumulation rate (Appleby and Oldfield, 1992, p. 746). The excess 210Pb activity in calcite as a function of time (t) is given by the formula:

Excess

2084

10

3

10

2

10

1

10

0

10

-1

10

-2

SPELEOTHEMS: Baskaran and Iliffe, 1993 Tanahara et al., 1998

10

-3

10

-4

10

-5

10

-6

Q

=

1 0.

dp m

1

0.

0.

Paulsen et al., 2003

10

×c

m

00

01

-2

×y

r

-1

1

YUCCA MOUNTAIN UZ CALCITE

10

(this work)

-8

10

-7

10

-6

10

-5

10

-4

10

-3

10

-2

10

-1

10

0

10

1

Growth Rate (mm/y) Fig. 9. (A) Evolution of excess 206Pb concentration produced from excess 210Pb (daughter isotope of 222Rn) with time in calcite grown in Rn-enriched environment. This concentration is a function of a mineral growth rate (growth curves from 101 to 107 mm y1 are shown) assuming a constant 210Pb influx Q. (B) Relations between mineral growth rates and concentrations of excess 206Pb in Yucca Mountain unsaturated zone calcite and speleothem calcite (calculated from published concentrations of excess 210Pb). Diagonal lines are loci of calculated excess 206Pb concentrations for different values of 210Pb fluxes (Q).

5.2.4. Distribution of excess 206Pb and 208Pb: the role of Ra absorbed on Mn-oxides Concentrations of excess 206Pb and 208Pb are highly variable, which is difficult to reconcile with remarkably uniform concentrations of parent U and Th in the TSw rocks (Bush et al., 1983; Zielinski et al., 1986; Gascoyne et al., 2002). In fact the Rn emanation will depend on concentrations of 226Ra and 224Ra (Semkow and Parekh, 1990; Morawska and Phillips, 1993), which are immediate Rn progenitors that may be heterogeneously distributed on cavity surfaces. Ra isotopes do not migrate far from the point of injection in oxidizing ground waters of moderate to low salinity and near neutral pH; they rapidly establish adsorption equilibrium between surface and solution (Krishnaswami et al., 1982; Copenhaver et al., 1993). It was shown that Ra is effectively extracted from waters by adsorption on MnO2-coated acrylic fibers (Moore, 1976; Michel et al., 1981; Krishnaswami et al., 1982; Moise et al., 2000). Eikenberg et al. (2001b) showed that MnO2coated discs extract almost 100% Ra from water at near-

Natural radionuclide mobility and U–Th–Pb dating of secondary minerals

neutral pH after 85 h of exposure. Therefore, Mn-oxides in the Yucca Mountain UZ may effectively scavenge Ra from slowly percolating near-neutral pH, dilute solutions. A presence in Mn-oxides of appreciable excesses of radiogenic Pb isotopes would indicate the possibility of effective Ra sorption, but, unlike in Pb-poor calcite, these excesses are very difficult to detect in Pb-rich Mn-oxides. However, 206Pb/204Pb ratios measured in Mn-oxides (18.13–18.60, Table 1) are slightly larger than the ratio of 206 Pb/204Pb = 18.127 ± 0.012 measured in the host tuff (Neymark et al., 1995). This difference cannot be explained by in situ 238U decay because the 238U/204Pb ratios in Mnoxides are smaller than in the host tuff (Fig. 4B). The largest 206 Pb excess of 1.29 ± 0.06 lg g1 is in Fe–Mn-oxide subsample HD2089APb3-Mn (calculated from data in Table 1), which is associated with calcite also having the largest measured 206Pb/204Pb ratios (Table 1). In addition, this Mn-oxide subsample has the largest concentration of Ba (>1%, Table EA1). Large specific surface areas of Mn-oxides (Fig. 3) favor sorption of Ra atoms, which will reside on the mineral surface. A large portion of Rn isotopes produced by decay of these adsorbed Ra atoms will be lost from Mn-oxides via a-recoil and diffusion. Therefore, air and water in the cavities with Mn-oxides may contain elevated concentrations of Rn compared to cavities without these minerals. As a consequence, calcite from cavities with Mn-oxides may accumulate elevated concentrations of Rn decay products. The much larger excess 206Pb relative to excess 208Pb, and virtual absence of excess 207Pb in calcite is in general agreement with the differences in half-lives of the parent isotopes of Rn. The presence of excess 208Pb and variable 208Pb/204Pb ratios in the UZ calcite may indicate that the proposed usage of 208Pb as an initial Pb proxy for the calcite U–Pb age calculations (Cliff and Richards, 2000; Cliff, 2001; Nemchin et al., 2006) should be done with caution in environments similar to the Yucca Mountain UZ. 5.3. Timing of flow and reliability of U–Pb dating New 207Pb/235U ages of opal/chalcedony ranging from 0.1329 ± 0.0080 to 9.10 ± 0.21 Ma (Table 3) support previous interpretations that the deposition of secondary cavity minerals from percolating water continued throughout the history of the mountain (Neymark et al., 2000, 2002). Slow long-term average growth rates of about 2.0 and 1.2 mm my1 obtained for coatings HD2055 and HD2233 (Section 4.4.2 and Fig. 8B) are in good agreement with previously published growth rates (Neymark et al., 2002; Paces et al., 2004). Measurable excesses of 207Pb were not found in U-poor UZ calcite except where it was closely associated with the U-enriched opal/chalcedony (Fig. 6). Only one calcite subsample not associated with opal contains a small 207Pb excess, which is easy to recognize due to extreme excesses of 206 Pb and 208Pb in this subsample. In most cases, therefore, very short-lived intermediate daughter Ra and Rn isotopes in the 235U decay chain do not contribute appreciably to the measured 207Pb abundances and cannot cause abnormally old 207Pb/235U ages in the Yucca Mountain UZ environ-

2085

ment. Recoil- and diffusion-related loss of intermediate decay products of 235U from opal, causing abnormally young 207 Pb/235U ages, however, may be an issue. 207Pb/235U ages of opal/chalcedony are in good agreement with the microstratigraphy of coatings (Fig. 8B) and define slow long-term average growth rates that are very similar to the growth rates based on 230Th–234U–238U opal ages (Paces et al., 2004). Because 234U and 230Th precede Ra and Rn in the 238 U decay chain, the U-series ages are not affected by potential additions of Ra and Rn decay products. New results presented in this paper support conclusions by Neymark et al. (2003) that 207Pb/235U ages of opal/chalcedony from the Yucca Mountain UZ are reliable estimates of the mineral formation ages and can be used for establishing the history of fracture flow through the site. 6. CONCLUSIONS Chemical and isotopic data for calcite, opal/chalcedony, and Mn-oxides in cavity coatings deposited by percolating water were used to evaluate the mobility of natural radionuclides in the unsaturated zone at Yucca Mountain and its effect on U-series and U–Th–Pb dating of these minerals. Isotopic compositions of Pb in secondary minerals record the time-integrated history of accumulation and differential mobility of U, Th, and their decay products. A conceptual model of processes occurring within a lithophysal cavity with growing secondary hydrogenic minerals is illustrated by Fig. 10. Samples of cavity-coating calcite contain excesses of 206 Pb, 207Pb, and 208Pb unsupported by in situ U and Th decay. Concentrations of excess 206Pb and 208Pb are highly variable, in contrast with remarkably uniform concentrations of parent U and Th in the TSw rocks. Mn-oxides can effectively adsorb Ra, which will reside on the mineral surface. A large portion of daughter Rn isotopes produced by decay of this adsorbed Ra will be lost from Mn-oxides via a-recoil and diffusion. Therefore, air and water in the cavities with Mn-oxides (Fig. 10A) may contain elevated concentrations of Rn compared to cavities without these minerals. As a consequence, calcite slowly growing in cavities with Mn-oxides may accumulate elevated concentrations of Rn decay products. 206 Pb/207Pb ratios measured in calcite vary over a wide range and reach a value of about 1400, higher than ever reported for natural Pb. Excess radiogenic Pb unsupported by in situ U and Th decay originates from decay of intermediate radioactive daughter isotopes of U and Th incorporated into the mineral during and after its formation. Excesses of 206Pb and 207Pb are found in U-poor calcite intimately intergrown with opal and are interpreted to be a result of a-recoil from U-rich opal (Fig. 10B). Therefore, 206 Pb/238U and 207Pb/235U ages of U-poor minerals closely associated with fine grains of U-enriched minerals may be grossly overestimated and should be interpreted with caution. Calcite from coatings devoid of opal contains excesses of 206Pb and 208Pb, but not appreciable excesses of 207Pb. These data indicate an extreme decoupling of 238U and 232 Th decay chains from the 235U decay chain. Most apparent 206Pb/238U and 208Pb/232Th ages of calcite are unrea-

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L.A. Neymark, Y.V. Amelin / Geochimica et Cosmochimica Acta 72 (2008) 2067–2089

Leaching of U, Th, and decay products from host tuff ter wa e t m i fro lc U to ca

Host Tuff:

222

Rn decay products to calcite r to

Film of Water U

rom

f

Opal

Vapor phase mineralization

Rn

te wa

Calcite

Opal 2 U ~ 10 μg/g

al op

U ~ 4 μg/g Rn ~ 124,000 Bq/m 3

Rn Rn h, Sr, f U, T oxide o n io nrpt Abso d Ra on M n Ba, a Rn

Recoiled nuclei Mn-oxide

Calcite -2 U ~ 10 μg/g Calcite volume affected by recoil implantation Fig. 10. Cartoons schematically showing processes of host rock dissolution, mineral precipitation, adsorption, and a-recoil causing mobility of U, Th, and their decay products within a lythophysal cavity with secondary hydrogenic minerals slowly growing from a film of percolating water (A) and a-recoil of U decay products from U-rich opal to U-poor calcite (B). Host tuff 222Rn concentration is calculated assuming equilibrium with parent 238U.

sonable, much older than the 12.8-Ma age of the host tuffs. Comparison of excess 206Pb in subsamples of cavity-coating calcite in the Yucca Mountain UZ and published data for speleothem calcite indicates that Rn emanation from the rock mass and U-enriched cavity-coating minerals (opal/ chalcedony and Mn-oxides) may cause the observed Pb isotope anomalies in slow-growing (mm my1) U-poor calcite from the Yucca Mountain UZ. The Rn emanation should not affect 206Pb–238U ages of fast-growing (mm y1) U-enriched speleothem calcites, but may affect dating of slowgrowing U-poor calcites. The presence of excess intermediate daughter isotopes does not appreciably affect 207Pb/235U ages of U-enriched opal/chalcedony, which are interpreted as mineral formation ages. New 207Pb/235U ages of opal/chalcedony range from 0.1329 ± 0.0080 to 9.10 ± 0.21 Ma and agree with previous data (Neymark et al., 2000, 2002) indicating that the deposition of secondary minerals started shortly after the eruption of the 12.8-Ma host tuff and continued throughout the history of the mountain. New estimates of slow long-term average growth rates of about 2.0 and 1.2 mm my1 are in good agreement with the published data. These results support previous conclusions that 207Pb/235U ages of opal/chalcedony from the Yucca Mountain UZ are reliable estimates of the mineral formation ages and can be used as geochronometers for establishing the history of fracture flow through the site.

ACKNOWLEDGMENTS The U.S. Geological Survey conducted this study in cooperation with the U.S. Department of Energy under Interagency Agreement DE-AI28-02RW12167. Loretta Kwak and Kiyoto Futa conducted ICP-MS concentration measurements and Sr isotope analyses of Mn-oxides, respectively. Reviews by D. Davis, K. Maher, R. Cliff, F. Corfu, and an unknown journal reviewer substantially improved the quality of this paper and are greatly appreciated. The guidance of the AE C. Johnson helped to streamline the paper. Any use of trade, product, or firm names in this paper is for descriptive purposes only and does not imply endorsement by the U.S. Government. The chemical analyses of secondary minerals and discussion of uranium sorption by Mn-oxides are presented in the Electronic Annexes EA-1 and EA-2, respectively.

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