U-series disequilibria in volcanic rocks from the Canary Islands: Plume versus lithospheric melting

Geochimica et Cosmochimica Acta, Vol. 67, No. 21, pp. 4153– 4177, 2003 Copyright © 2003 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/03 $30.00 ⫹ .00

Pergamon

doi:10.1016/S0016-7037(03)00308-9

U-series disequilibria in volcanic rocks from the Canary Islands: Plume versus lithospheric melting CRAIG C. LUNDSTROM,1,* KAJ HOERNLE,2 and JIM GILL3 1

Department of Geology, University of Illinois at Urbana Champaign, 1301 W. Green St., Urbana, IL 61801, USA Department of Volcanology and Petrology, GEOMAR Research Center, Wischhofstr. 1-3, 24148 Kiel, Germany 3 Earth Sciences Dept., University of California Santa Cruz, Santa Cruz, CA 95064, USA

2

(Received March 15, 2002; accepted in revised form May 2, 2003)

Abstract—U-series disequilibria are presented for Holocene samples from the Canary Islands and interpreted with special emphasis on the separate roles of plume vs. lithospheric melting processes. We report Th and U concentrations and (238U)/(232Th), (230Th)/(232Th), (230Th)/(238U) and (234U)/(238U) for 43 samples, most of which are minimally differentiated, along with (226Ra)/(230Th) and (231Pa)/(235U) for a subset of these samples, measured by thermal ionization mass spectrometry (TIMS). Th and U concentrations range between 2 and 20 ppm and 0.5 and 6 ppm, respectively. Initial (230Th)/(238U) ranges from 1.1 to 1.6. (226Ra)/(230Th)o ranges between 0.9 and 1.8 while (231Pa)/(235U)o ranges between 1.0 and 2.0. Our interpretation of results is based on a three-fold division of samples as a function of incompatible element ratio, such as Nb/U. The majority of samples have Nb/U ⫽ 47 ⫾ 10, similar to most MORB and OIB. Higher ratios are found exclusively in alkali basalts and tholeiites from the eastern Canary Islands whereas lower ratios are exclusively found in differentiated rocks from the western Canary Islands. Those with ordinary Nb/U ratios are attributed to melting within the slowly ascending HIMU-dominated Canary plume. Higher Nb/U, generally found in more silica rich basalts from the eastern islands, is attributed to lithospheric contamination. Based on their trace element characteristics, two possible contaminants are amphibole veins (⫾ other minerals) crystallized in the mantle from previous plume-derived basanite or re-melted plume-derived intrusive rocks. The high Nb/U signature of these materials is imparted on a melt of the lithosphere created either by the diffusive infiltration of alkalis or by direct reaction between basanites and peridotite. Mixing between plume-derived basanite and lithospheric melt accounts for the U-series systematics of most eastern island magmas including the well-known Timanfaya eruption. Lower Nb/U ratios in differentiated rocks from the western islands are attributed to fractional crystallization of amphibole ⫾ phlogopite ⫾ sphene from basanite during its ascent through the lithosphere. Based on changes in disequilibria, phonolites and tephrites are interpreted to result from rapid differentiation of primitive parents within millennia. Copyright © 2003 Elsevier Ltd U-series disequilibria can differ between plume melts and lithospheric melts (Bourdon et al., 1998). Indeed, basalts derived from continental lithosphere have high (231Pa)/(235U) indicating that large parent-daughter disequilibria can be generated during melting of continental lithosphere (Asmerom et al., 2000). Although Sr-Nd-Pb-O-Os isotope studies have demonstrated that lithospheric contamination affects the geochemistry of some Canary Island basaltic magmas (e.g., Hoernle et al., 1991a; Hoernle and Tilton, 1991; Hoernle, 1998; Thirlwall et al., 1997; Widom et al., 1999), none of the previous U-series studies considered lithospheric effects in their interpretations. For instance, Sigmarsson et al. (1998) suggested that the Timanfaya eruption on Lanzarote was the product of mixed melts derived from garnet pyroxenite and garnet peridotite within the upwelling plume. Thomas et al. (1999) also identified the importance of mixing in this eruption but attributed the end member melts to various degrees of melting of homogeneous peridotite at slightly different depths in the plume. Finally, decreases in 230Th-238U disequilibria for a basanite to phonolite sequence from Tenerife were attributed solely to time of differentiation (Hawkesworth et al., 2000). Here we investigate the Holocene volcanic rocks of the entire Canary Islands, an end member in the spectrum of plumes, to separate the effects of melting in the oceanic lithosphere vs. in the plume. Ocean

1. INTRODUCTION

Despite a rapid increase in the number of high precision geochemical data, major questions regarding the mechanisms of hot spot melting in the oceanic upper mantle remain unanswered. The issues include determining the relative contributions of melt from plume vs. lithospheric sources and constraining the rates and time scales of magma production and ascent. A powerful tool for examining this latter issue is measurement of disequilibria between short-lived members of the uranium decay series (U-series disequilibria). Because such data provide the ability to probe the time scale of processes in the 0- to 350-ka range, U-series data allow quantification of the rates of solid mantle upwelling, melt ascent and crystal fractionation. A number of U-series studies have examined ocean island basalt genesis, primarily focusing on the generation of disequilibria during upwelling of plume material (Cohen and O’Nions, 1993; Chabaux and Allegre, 1994; Turner et al., 1997; Bourdon et al., 1998; Sigmarsson et al., 1998; Sims and DePaolo, 1997; Thomas et al., 1999; Pietruszka et al., 2001; Kokfelt et al., in press). However, lithospheric contributions to ocean island magmas can also be significant (Class and Goldstein, 1997) and * Author to whom ([email protected]).

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Fig. 1. The Canary and Selvagen Islands together with the seamounts to the northeast form the ⬃70-Ma Canary Hotspot Track (modified from Geldmacher et al., 2001). Oldest available ages are given in parenthesis (in Ma). The greater than sign (⬎) is used when the oldest dated samples belong to the posterosional volcanism based on their geochemical composition (206Pb/204Pb ⬍ 19.5). Shaded field shows the extent of a seismic chaotic facies (UCF), which is interpreted to be of volcanic origin and to decrease in age southwards from late Cretaceous to early Tertiary (Holik et al., 1991).

island basalts typically range in composition from silica undersaturated nephelinites, basanites and alkali basalts (and associated differentiates) to tholeiites. In the Canary Islands, most basalts are alkalic rather than tholeiitic, even during the shieldbuilding stage. The most tholeiitic rocks have erupted only on the eastern islands and only during the Pliocene and Quaternary. Both globally and in the Canary Islands, sorting out the genesis of alkalic and tholeiitic basalts demands a comprehensive approach incorporating major and trace elements along with radiogenic and short-lived isotopes. We use several techniques to distinguish the processes leading to the diversity of magmas erupted within the Canary Islands. We show that three separate trends in silica enrichment occur, distinguished by trace element systematics and U-series disequilibria. One trend leading to slightly silica under-saturated alkali basalts can be attributed to decreasing pressure and increasing melting column length within the plume. A second trend of silica enrichment leading to tholeiites can be attributed to the melting of lithosphere that has been metasomatized recently by plume magmas. Lastly, we suggest that alkali basalt differentiation primarily occurs by melt-rock reaction in cold lithosphere. We conclude that the repeated passage of alkalic plume magmas through the lithosphere has led to extensive modification and subsequent melting of thick, cold oceanic lithospheric mantle.

2. GEOLOGIC CONTEXT OF SAMPLES

The Canary Islands are located between 100 and 500 km off the western coast of Africa at 28°N latitude. The archipelago consists of seven main islands, which show a crude east (Fuerteventura, ⬃24 Ma) to west (El Hierro, ⬃1.1 Ma) age progression in subaerial volcanism (Fig. 1). In this paper we will refer to La Palma, La Gomera, Hierro, and Tenerife as “western islands,” vs. Gran Canaria, Fuenteventura and Lanzarote, referred to as “eastern islands.” Together with the Selvagen Islands and a group of seamounts to the northeast, the Canary Islands form an ⬃800 km long and 450 km wide volcanic belt which decreases in age from the northeast (⬃68 Ma, Lars Seamount) to the southwest. This is interpreted to represent the Canary hotspot track (Holik et al., 1991; Geldmacher et al., 2001). This hotspot track can be considered a global end member in several respects. First, the islands are located on the oldest oceanic crust of all ocean island groups, with lithospheric age ranging from ⬃150 Ma in the west to ⬃180 Ma in the east (Roest et al., 1992; Hoernle, 1998; Schmincke et al., 1998). Although oceanic lithosphere older than ⬃70 Ma is generally considered to have a thickness of 100 to 125 km (e.g., Parsons and Sclater, 1977), studies of mantle xenoliths from the east-

U-series disequilibria in the Canary Islands

ernmost island of Lanzarote show an anomalously high geothermal gradient with temperatures of 1100°C at ⬃8 kbar (Neumann et al., 1995). The high geotherm is interpreted to reflect thermal erosion of the base of the lithosphere beneath the eastern Canary Islands, probably reflecting long-term thermal exposure to the Canary plume. Since the younger, western islands have only recently been located above the hottest part of the plume, the thickness of the lithosphere beneath these islands will approach that based on the plate’s conductive cooling profile. Second, due to the very slow plate motion (⬃12 mm/yr; Geldmacher et al., 2001), the Canary hotspot and lithosphere have been interacting for tens of millions of years beneath the older volcanoes (e.g., eastern islands). The resulting metasomatism of the sub-Canarian lithosphere (crust and mantle) ranges from cryptic to crystallization of exotic phases (i.e., amphibole, phlogopite, apatite, zircon, ilmenite and pyrochlore) to the formation of amphibole cumulates and carbonatite veins. Finally, the buoyancy flux of the Canary hotspot has been estimated at 1.0 Mg/s (Sleep, 1990), placing the Canary plume among the weakest plumes. This aspect suggests that the solid upwelling and, therefore, melting rates beneath the Canaries should be among the lowest of all hot spots. Together, the above features make the Canaries an excellent location to investigate the rates of melt generation and lithospheric interaction using U-series disequilibria. The geochemistry of Canary Island volcanic rocks is now well understood in a volcano evolutionary context. The older Canary Island volcanoes show multiple growth cycles, referred to here as shield-building (both seamount and subaerial) and posterosional. During the shield-building cycle, volcanic rocks have HIMU-type Sr-Nd-Pb isotopic compositions (Hoernle et al., 1991a; Hoernle and Tilton, 1991; Marcantonio et al., 1995; Simonsen et al., 2000), which probably reflect a plume component dominated by ⬍ 2-Ga recycled oceanic crust (Thirlwall et al., 1997; Widom et al., 1999). Rocks of the posterosional cycle on the eastern islands are characterized by less radiogenic Sr-Pb and more radiogenic Nd, interpreted to reflect the involvement of depleted midocean ridge basalt source mantle (DMM) also in the plume (Hoernle et al., 1991a; Hoernle and Tilton, 1991). In both growth cycles, nephelinites and basanites dominate the early stage, more silica-saturated basaltic melts are produced at the peak of the cycle, and evolved alkalic melts with varying abundances of alkali basalts through nephelinites appear during the waning stages (Hoernle and Schmincke, 1993a, 1993b). In contrast to some other hotspots such as Hawaii, tholeiites are relatively uncommon, having been erupted primarily on the eastern Canary Islands (Ibarrola, 1970; Fuster, 1975). In both growth cycles, the more silica saturated mafic and evolved volcanic rocks generally have more radiogenic Sr and less radiogenic Nd and Pb (more enriched or EM-type) isotopic compositions than the associated basanites. These more enriched isotopic compositions are interpreted to result from lithospheric interaction (Hoernle et al., 1991a; Hoernle, 1998; Widom et al., 1999). Further information for each Canary Island is given in Appendix A. Holocene volcanism has occurred on all of the islands except La Gomera (last active in the Pliocene) with historic activity occurring on four islands across the entire chain (El Hierro, La Palma, Tenerife and Lanzarote). The youngest samples from each island belong

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to different stages in volcano evolution discussed above including a range of lithospheric interactions. 3. SAMPLES AND ANALYTICAL METHODS Forty-three lavas were studied for U-series disequilibria and are representative of the youngest available from six Canary Islands (Table 1). All are hand-picked interior pieces from the youngest lavas on their island. Twenty-five are from historical 14C- or Ar/Ar-dated eruptions. The remainder are estimated to be 3 to 25 ka old based on field evidence and experience; estimated ages are indicated by “?” in Table 2 and are probably uncertain by a factor of 2. The subscript “o” throughout this paper indicates the radioactive disequilibrium at the time of eruption using the age given for this correction. Further information about the samples is given in Appendix A. For simplicity we refer to all of them as “Holocene” to distinguish them from older samples in the figures. They are representative in major and trace elements of those erupted on their island throughout the Quaternary. Major and trace element concentrations were measured by XRF (GEOMAR) and ICPMS (University of Kiel). Table 1 contains results for standards. Sr, Nd, Pb, Hf, O and He isotope ratios have been measured for most of the samples and will be published separately. Os isotope ratios for about half of our samples were published and discussed by Widom et al. (1999). Th and U concentrations and isotope ratios were determined by thermal ionization mass spectrometry (TIMS) at the University of California Santa Cruz. Sample powders (100 – 400 mg) were spiked with 229Th and 233U, digested in concentrated HF (5–10 mL) and HClO4 (0.2 mL) and refluxed on a hotplate at 130 to 150°C for 5 d. After gentle drying, the sample was refluxed in 6 N HCl for 24 h (130 –150°C) to gain complete solution. The entire procedure was repeated if dissolution was incomplete (⬃10% of the dissolutions). After gentle drying, the sample was dissolved in 8 N HNO3 for U and Th separation by anion chromatography. All reagents used during the dissolution and chemical separation process were Seastar or Optima brand and blanks for U and Th were 10 and 15 pg, respectively. Procedures for separations and mass spectrometry of Ra and Pa are those given in Lundstrom et al. (1998). Blanks for Ra and Pa were below detection limits. Initial isotope dilution analyses for U and Th not reported here were performed using concentrated HNO3 (2–3 mL) instead of HClO4 during the dissolution. Although the U concentrations measured using this procedure showed the same level of reproducibility as dissolutions done later with HClO4, 90% of the Th concentrations were ⬎ 2% higher than the dissolutions done using HClO4. We attribute these higher Th concentration values to the lack of spike/sample equilibration due to the incomplete removal of fluoride ion during dissolution without HClO4. Based on our cumulative experience in the UCSC laboratory, spike-sample equilibration for both Th and Pa critically depend on inclusion of HClO4 during the dissolution process. U was loaded on Re filaments with graphite and analyzed on the mass spectrometer by peak jumping between 233, 234 and 235. Th was loaded on triple Re filaments with boric acid. All Re used was four pass zone refined (H. Cross). U, Th and most Ra measurements were performed on a VG 54-30 mass spectrometer having a 30-cm energy filter and equipped with an ion counting Daly detection system having a 30-cpm dark current. Dead time on this detector is 20 ns (verified by repeat analysis of NBS U010). All Pa measurements and some Ra measurements were made on a NBS 1260 equipped with an electron multiplier having a 12-ns deadtime and a 12-cpm dark current. 4. RESULTS

4.1. Geochemical Background Samples range from “primitive” basanites, alkali basalts, and tholeiites to evolved trachytes and tephriphonolites (Table 1). By primitive we mean minimally differentiated rocks with Mg# (⫽MgO/[MgO ⫹ 0.8 FeOT]) ⬎ 60, ⬎ 10% MgO, ⬎ 200 ppm Ni, and ⬎ 300 ppm Cr. By this definition, 84% of the samples are primitive. Holocene basanites (SiO2 of 41– 45 wt.%) occur

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C. C. Lundstrom, K. Hoernle, and J. Gill Table 1. Major and trace element concentrations for U-series samples.

Sample #

Age (yr)

Lanzarote EL1 AD 1824 EL10 AD 1730 EL11 AD 1730 EL5 AD 1731 EL3 AD 1731 EL13 AD 1732 EL8 AD 1736 EL15 20,000 EL17 20,000 EL18 20,000 EL21 20,000 EL20 20,000 EL23 25,000? Fuerteventura EF5 5000? EF1 5000? KFS240 5000? EF10 5000? EF7 5000? EF9 5000? EF12 5000? EF6 5000? Gran Canaria EGC7 3075 BP EGC3 3080 BP EGC6 3075 BP EGC8 10,000? Tenerife ET2 25,000? ET7 25,000? ET6 AD 1705 ET8 25,000? ET1577 AD 1909 ET28 800 BP La Palma LP261925A 5000? LP71-7L AD 1971 LP124794 AD 1949 LP124793.1 AD 1949 ELP7 10,000? LP43e AD 1646 ELP1 AD 1480 WHJLP.92.11 2310 BP Hierro EH8 1793? AD EH11 5000? EH4 2900 BP EH17 25,000? Standards n⫽6 BHV0-1 QLO-1 DR-N

Mg#

SiO2

TiO2

Al2O3

FeO1

MnO

MgO

CaO

Na2O

K2 O

P2O5 H2O* CO2* Total*

72.5 73.1 72.6 68.6 68.0 68.7 66.8 67.0 67.7 68.9 69.7 69.6 66.6

44.72 43.00 43.08 47.49 49.76 48.34 50.75 47.97 48.34 46.62 46.07 44.54 45.72

2.65 2.98 2.97 2.49 2.51 2.60 2.44 2.48 2.60 2.68 2.72 3.08 3.03

11.92 11.53 11.56 13.28 13.20 12.71 13.47 13.22 13.13 12.90 12.42 12.41 12.98

10.96 11.76 11.89 11.23 10.71 11.21 10.50 11.12 10.88 11.24 11.38 11.97 11.59

0.18 0.20 0.20 0.17 0.16 0.16 0.15 0.17 0.17 0.18 0.18 0.19 0.18

12.94 14.31 14.11 11.03 10.20 11.06 9.47 10.12 10.24 11.17 11.75 12.30 10.38

11.18 10.05 9.98 9.67 9.29 9.51 9.17 10.08 10.03 10.23 10.44 10.24 10.76

3.37 3.63 3.69 3.25 3.10 3.10 3.08 3.38 3.24 3.35 3.24 3.29 3.49

1.16 1.52 1.51 0.83 0.71 0.92 0.63 0.83 0.83 0.99 1.05 1.26 1.14

0.91 1.00 0.99 0.54 0.36 0.40 0.34 0.61 0.54 0.64 0.72 0.72 0.73

0.11 0.13 0.16 0.16 0.27 0.18 0.09 0.15 0.10 0.13 0.24 0.37 0.15

0.03 0.03 0.04 0.03 0.06 0.02 0.02 0.04 0.02 0.05 0.02 0.03 0.02

100.04 100.33 100.15 100.90 100.06 100.84 100.99 100.73 100.44 99.38 100.31 100.20 100.63

71.2 69.1 69.5 68.0 66.8 66.1 65.6 62.2

43.30 45.93 41.94 45.07 45.68 49.64 49.32 49.34

3.16 2.84 3.38 3.29 2.87 2.43 2.48 2.91

12.05 12.81 11.63 12.17 13.09 13.46 13.50 13.69

12.27 11.56 12.78 11.94 12.03 10.92 11.25 11.33

0.18 0.18 0.20 0.18 0.18 0.16 0.16 0.17

13.62 11.58 13.09 11.40 10.84 9.56 9.62 8.38

10.25 10.51 11.19 10.82 10.12 9.39 9.23 9.51

3.21 2.98 3.28 3.18 3.38 3.33 3.32 3.36

1.34 1.11 1.48 1.28 1.13 0.68 0.66 0.86

0.61 0.50 1.01 0.65 0.65 0.41 0.43 0.44

0.21 0.41 0.35 0.28 0.30 0.21 0.19 0.15

0.03 0.04 0.10 0.03 0.02 0.03 0.04 0.03

100.11 100.56 100.40 100.95 100.52 100.05 100.82 100.59

72.4 68.4 68.2 67.2

42.68 42.07 43.70 43.59

3.72 4.04 3.68 3.93

11.51 12.27 13.29 12.06

11.43 11.44 10.74 12.23

0.17 0.19 0.19 0.19

13.44 11.13 10.34 11.24

11.93 13.08 10.90 11.35

2.89 3.28 4.22 3.08

1.49 1.51 2.20 1.51

0.72 0.97 0.73 0.80

0.67 0.18 0.16 0.37

0.03 0.04 0.03 0.03

100.47 100.19 100.37 100.37

65.3 64.7 62.3 61.8 50.3 43.8

44.73 46.18 43.99 47.51 44.90 56.28

3.28 2.93 3.67 3.05 3.89 1.79

13.60 13.64 13.95 14.82 16.69 18.16

11.88 11.46 12.59 11.09 11.62 5.78

0.19 0.18 0.19 0.18 0.19 0.17

10.02 9.41 9.33 8.04 5.28 2.02

10.34 9.73 10.63 9.16 10.56 3.94

3.68 4.13 3.53 3.95 4.26 7.26

1.54 1.71 1.39 1.40 1.69 4.02

0.72 0.60 0.72 0.79 0.91 0.56

0.14 0.23 0.13 0.20 0.16 0.16

0.03 0.03 0.04 0.03 0.04 0.03

100.13 100.37 100.09 100.41 100.89 99.83

68.4 59.1 60.0 52.4 62.5 60.3 58.3 43.7

43.89 43.83 45.12 46.50 43.29 43.45 44.17 50.41

3.48 3.88 3.43 3.48 3.88 3.83 3.88 2.38

11.66 13.89 14.00 16.20 13.20 13.34 14.25 18.51

12.38 12.72 12.49 10.76 12.42 12.30 12.12 7.66

0.20 0.21 0.20 0.21 0.20 0.20 0.22 0.21

12.06 8.24 8.41 5.32 9.31 8.38 7.59 2.67

11.13 11.12 10.73 9.20 11.71 12.43 10.55 6.99

3.21 3.80 3.65 5.09 3.74 3.60 4.38 7.22

1.28 1.41 1.29 2.33 1.38 1.57 1.96 3.21

0.70 0.88 0.66 0.90 0.85 0.89 0.86 0.74

0.19 0.21 0.18 0.68 0.23 0.00 0.18 0.23

0.04 0.08 0.01 0.10 0.04 0.00 0.03 0.04

100.57 100.98 101.22 100.68 99.94 99.61 100.37 99.40

68.3 65.6 64.4 52.0

41.80 42.05 44.61 46.60

4.37 4.19 3.52 3.84

9.56 11.48 12.62 16.38

14.99 13.88 13.62 10.63

0.19 0.19 0.20 0.19

14.50 11.86 11.04 5.17

10.99 11.32 9.63 9.09

2.23 3.10 3.23 4.85

0.87 1.08 0.93 2.07

0.49 0.82 0.58 1.17

0.18 0.12 0.13 0.22

0.03 0.03 0.02 0.03

100.83 100.99 100.79 100.19

49.72 65.35 52.68

2.79 0.61 1.06

13.52 16.21 17.58

12.47 4.28 9.52

0.17 0.10 0.22

7.22 1.01 4.33

11.22 3.16 6.87

2.33 4.28 2.96

0.52 3.61 1.69

0.27 0.26 0.23

0.10 2.26

100.23 98.96 99.40

Samples from historical eruptions are given oldest to youngest; others are arranged by decreasing Mg# within islands. Question marks indicate age uncertainty when neither historical nor 14C dated; see text. Major elements are by XRF and normalized to an anyhdrous basis. Original H2O and CO2 contents and Totals are indicated by * and are given for reference only. Mg# is calculated assuming FeO/FeOT ⫽ 0.8. Trace elements Sc to Ba are by XRF, Li to Pb by ICPMS, and Th-U by TIMS-ID. Results for international standards were obtained along with unknowns.

on all the islands, but alkali basalts and tholeiites (45–51%) are only from Tenerife and the two easternmost islands, Fuerteventura and Lanzarote, with tholeiites only from these latter two. The evolved volcanic rocks with Mg# ⬍ 60 are from Tenerife and La Palma. In general, the samples in this study are geochemically representative of our larger sample set discussed below.

The SiO2 content of primitive volcanic rocks with ⬍ 20% phenocrysts varies systematically with other major and trace elements. The Al2O3 content (10.6 –14.8 wt.%) correlates positively with SiO2, whereas FeOT (10.4 –14.1 wt.%), CaO (9.2– 13.1 wt.%), CaO/Al2O3 (0.6 –1.1 wt.%), MgO (8 –16 wt.%), TiO2, K2O, P2O5, MnO, V, Cr and incompatible trace element contents correlate inversely with SiO2. Notably, the basanitic

U-series disequilibria in the Canary Islands

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Table 1. continued Sample # Lanzarote EL1 EL10 EL11 EL5 EL3 EL13 EL8 EL15 EL17 EL21 EL20 EL20 EL23 Fuerteventura EF5 EF1 KFS240 EF10 EF7 EF9 EF12 EF6 Gran Canaria EGC7 EGC3 EGC6 EGC8 Tenerife ET2 ET7 ET6 ET8 ET1577 ET28 La Palma LP261925A LP71-7L LP124794 LP124793.1 ELP7 LP43e ELP1 WHJLP.92.11 Hierro EH8 EH11 EH4 EH17 Standards BHV0-1 QLO-1 DR-N

Sc

V

Cr

19 16 20 20 14 14 23 19 20 17 13 22 18

228 239 241 210 196 201 187 203 209 218 229 243 239

547 565 580 418 417 468 391 384 420 439 486 456 372

20 21

262 253

23 20 24 16 22

Co

Ni

Cu

Zn

Ga

Rb

Sr

Y

Zr

Nb

Ba

Li

63 65 63 60 57 60 52 52 55 58 60 65 56

346 440 430 314 299 335 279 265 255 297 309 322 229

84 59 67 87 77 62 83 74 60 100 65 63 69

101 125 111 118 107 104 113 113 120 109 103 105 108

18 18 18 19 20 15 18 19 19 20 16 16 19

20 32 31 12 11 15 11 11 12 19 17 21 20

1023 1015 1018 621 438 470 441 758 702 754 825 784 771

31 29 30 27 20 24 22 27 23 27 28 25 27

268 300 309 215 182 220 182 191 201 212 221 246 248

70 75 84 46 32 41 33 51 44 49 57 65 61

515 534 544 309 210 257 194 358 348 417 436 414 419

11.3 7.9 5.0 6.4 4.0 3.9 5.4

64 62

89 52 80 63 67 81 85 80

124 106 107 120 123 131 135 127

19 17 15 21 22

27 18 32 19 19 10 7 10

666 565 929 722 690 561 528 524

26 22 29 27 27 25 29 27

226 199 285 263 247 189 194 245

60 36 82 60 56 33 36 46

398 289 562 369 345 228 232 222

5.6

56 53 52 54 51

395 286 329 254 260 248 264 187

17 16

279 238 186 192 226

491 457 425 466 363 362 367 316

27 30 22 19

314 354 322 296

759 497 581 449

66 60 49 61

296 183 189 234

85 104 79 75

94 102 118 119

12 17 19 19

24 23 51 30

896 1187 1337 1117

25 32 28 30

257 284 397 291

55 71 87 62

505 568 807 566

19 26 22 14 11 8

279 268 305 237 292 92

399 381 313 238 ⬍18 ⬍18

55 58 61 51 38 9.8

221 217 188 224 18 5

67 88 60 95 37 15

118 123 127 122 121 112

17 20 21 18 24 24

32 40 26 22 27 84

855 676 886 893 1114 427

30 29 27 30 31 29

288 309 263 277 296 481

73 78 62 71 85 109

479 853 417 496 540 860

28 24 23 11 29

315 303 300 267 351

726 224 331 71 362

299 121 131 47 168

100 91 76 61 100

105 120 133 143 116

14 20 19 20 18

29 23 27 50 36

895 1055 931 1389 1031

27 34 30 37 31

251 309 257 383 297

62 68 61 102 77

472 429 412 748 554

16 6

315 147

239 ⬍18

53

113 8

83 37

134 138

19 25

39 88

1159 1745

36 37

364 543

88 145

590 1023

9.3

27 23 23 1

368 334 281 233

700 491 369 23

85 76 65 34

404 312 323 42

85 98 83 69

130 131 142 155

17 16 21 25

13 19 16 42

594 867 705 1254

28 33 30 43

257 296 278 554

36 59 50 99

192 319 258 478

4.6 6.4 6.9

25 ⬍1 24

307 46 222

276 ⬍18 31

8.1 7.6 9.6

117 4 25

106 23 62

102 65 148

20 26 24

7 70 72

404 321 386

29 25 29

176 170 132

15 3 4

126 1357 357

3.3 8.7 5.2

48 60

through tholeiitic basalts erupted on Lanzarote between 1730 and 1736 (Timanfaya eruption) show similar correlations to the entire data set of primitive volcanic rocks from the Canaries but form much tighter linear correlations. When considering the whole data set (both primitive and differentiated samples), SiO2 increases with decreasing MgO along two separate differentiation or melting trends (Fig. 2). One trend corresponds to primitive rocks representing the basanite to tholeiite sequence observed on the eastern islands. A second trend contains alkalic series differentiates which remain low in SiO2 as MgO de-

7.0

6.3 6.1 6.2 5.6 8.4 6.0 7.3 6.4 7.9

11.0 5.0 7.2

creases to 5 wt.%, followed by an abrupt increase in SiO2 at lower MgO. The extension of this trend leads to syenites and phonolites such as those observed on Tenerife (Wolff et al., 2000). Primitive mafic volcanic rocks, as well as differentiates, show large variations in ratios of highly incompatible trace elements such as (K, Rb, Ba, Nb, Ta, Zr)/(U, Th, LREE). For example, mafic volcanic rocks have Nb/U ⫽ 34 to 111, K/U ⫽ 5000 to 17,000, and Zr/U ⫽ 100 to 600. These incompatible element ratios show inverse hyperbolic correlations as a func-

4158

C. C. Lundstrom, K. Hoernle, and J. Gill Table 1. continued

Cs

Ba

La

0.34 0.54 0.48 0.26 0.22 0.23 0.19

478 527 480 297 186 214 181

69.5 58.8 54.9 34.2 19.1 21.0 18.6

0.19 347

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Hf

Ta

Pb

Th

U

145 16.64 126 15.39 116 14.90 72 9.07 42 5.77 46 6.30 41 5.47

63.7 61.3 58.2 37.2 24.6 26.8 24.4

11.4 11.6 11.0 7.72 5.85 6.17 5.96

3.37 3.45 3.26 2.47 1.92 2.01 2.02

9.66 9.64 9.05 7.17 5.36 5.59 5.86

1.23 1.26 1.20 0.99 0.79 0.82 0.84

5.96 6.13 5.94 5.11 4.23 4.32 4.51

1.00 1.01 0.99 0.87 0.74 0.75 0.77

2.54 2.56 2.45 2.25 1.89 1.84 1.95

0.30 0.29 0.29 0.27 0.23 0.23 0.24

1.82 1.78 1.71 1.66 1.42 1.39 1.46

0.24 0.23 0.22 0.22 0.19 0.19 0.19

5.81 6.70 6.59 5.09 4.18 4.97 4.45

3.94 4.52 4.18 2.86 2.04 2.50 1.98

2.89 3.03 2.73 2.00 1.32 1.31 1.28

41.6

84 10.35

41.5

8.36 2.60

7.51 1.02

5.23

0.89

2.27

0.28

1.67

0.23

5.00

2.77

2.75

6.93 6.03 6.04 3.32 2.16 2.34 1.96 5.02 4.49 4.91 5.72 5.40 4.46

1.74 1.59 1.62 0.84 0.56 0.60 0.52 1.14 1.18 1.30 1.47 1.28 1.11

0.39 398

39.7

83 10.23

40.9

8.26 2.50

7.29 0.97

4.81

0.81

2.02

0.24

1.42

0.18

5.46

3.64

2.06

0.36 337

36.3

79

9.96

40.6

8.36 2.60

7.65 1.03

5.22

0.89

2.27

0.28

1.64

0.22

5.53

3.51

2.14

0.17 218 0.21 226

24.8 23.5

54 53

6.95 7.11

29.0 30.8

6.55 2.12 7.07 2.26

6.16 0.88 6.79 0.96

4.64 4.96

0.80 0.87

2.02 2.21

0.26 0.27

1.57 1.67

0.21 0.23

4.59 5.81

2.26 3.06

1.54 1.47

4.10 2.54 6.55 3.87 3.70 2.50 2.46 2.11

1.00 0.68 1.66 0.88 0.92 0.61 0.46 0.41

0.42 525 0.79 777 0.46 520

73.3 70.4 50.0

156 18.36 140 15.93 108 13.52

71.3 59.8 55.1

12.5 3.67 10.71 1.32 10.6 3.15 9.24 1.15 10.6 3.20 9.16 1.17

6.27 5.64 5.66

1.05 0.97 0.95

2.61 2.54 2.44

0.30 0.31 0.29

1.88 2.03 1.78

0.25 0.27 0.24

6.72 8.16 6.85

4.02 4.24 3.66

3.25 5.65 3.29

4.87 7.38 8.94 5.11

1.23 1.81 2.41 1.36

0.43 458

52.0

106 12.73

49.6

9.71 3.01

8.69 1.16

5.95

1.03

2.70

0.32

1.99

0.27

6.14

4.15

3.38

0.36 402 0.40 437

44.7 52.4

94 11.64 107 12.61

46.6 48.6

9.42 2.96 9.62 3.02

8.37 1.13 8.53 1.17

5.69 5.91

0.98 1.05

2.50 2.68

0.30 0.33

1.81 2.01

0.24 0.27

5.82 5.94

3.77 3.90

2.85 3.24

4.96 6.57 4.21 5.19

1.27 1.57 0.99 1.37

4.87 1.18 12.60 3.17 6.82 6.53 5.65 11.50 3.77 7.31 9.08 4.46 7.78 19.80

1.72 1.86 1.58 3.32 2.03 2.40 2.20 6.20

2.48 3.45 3.23

2.14 2.79 2.60

0.75 1.09 0.99 2.30

0.04 0.46 1.10

3.28 7.84 2.10

0.43 426 0.35 337

69.9 51.3

143 16.73 106 13.44

64.5 52.0

12.0 3.60 10.30 1.36 10.1 3.03 8.54 1.16

6.74 6.01

1.17 1.05

3.02 2.65

0.36 0.33

2.19 2.00

0.29 0.27

6.75 6.01

4.06 3.45

0.52 523

74.7

152 17.53

66.7

12.1 3.54 10.30 1.36

6.70

1.16

3.00

0.37

2.30

0.31

7.16

4.44

0.63 602

79.4

165 19.35

75.3

13.6 4.05 11.59 1.55

7.68

1.33

3.42

0.42

2.61

0.34

8.25

5.51

0.30 199 0.29 302 0.35 241

29.3 48.4 39.6

68 9.31 106 13.30 89 10.94

40.4 54.8 44.3

9.04 2.75 11.2 3.38 9.30 2.89

7.99 1.06 9.64 1.28 8.56 1.16

5.24 6.25 5.97

0.87 1.03 1.03

2.17 2.57 2.63

0.25 0.30 0.32

1.52 1.74 1.90

0.20 0.23 0.26

6.18 6.99 6.57

2.5 12.9 23.9

1.16 0.53 3.35 1.11 5.98 2.00

1.86 0.38 3.72 0.62 6.07 0.92

2.61 3.84 5.15

0.58 0.79 0.95

1.74 2.28 2.44

0.25 0.32 0.32

1.67 2.04 1.97

0.25 0.30 0.27

0.60 2.46 4.44

⬍0.05 7.2 0.95 160 0.10 136

0.6 2.0 10.5 22.8 15.6 37.7

0.40 3.01 5.29

tion of U and other incompatible element concentration regardless of the age of volcanism (Hoernle and Schmincke, 1993; Fig. 3). Basanites have lower ratios while tholeiites, with lower incompatible element contents, have higher ratios. In Figure 3, most samples with U ⬎ 1.25 ppm are basanites. Samples with high (K, Rb, Ba, Nb, Ta, Zr)/(La, U, Th) come almost exclusively from the easternmost Canary Islands of Fuerteventura and Lanzarote. Because these trends also exist with Th as denominator and are observed in Holocene lavas with (234U)/ (238U) ⫽ 1 (see below), alteration and mobilization of U cannot explain the high (Nb, Zr, K)/U (Fig. 3 inset). At the other

3.52 2.85

3.01 4.32 3.62 8.74

extreme, differentiated samples such as tephrites and phonotephrites from the western islands have low ratios. The relationship between rock type (MgO-SiO2) and Nb/U is readily apparent in Figure 2. Although Sr, Nd, and Pb isotope ratios do not correlate significantly with the U-series data, the basanites appear to be a mix of HIMU and DMM components, while tholeiites trend toward an EM component (see Widom et al., 1999). The range in isotopic composition is 87Sr/86Sr ⫽ 0.7029 to 33, 143Nd/ 144 Nd ⫽ 0.5129 to 30, and 206Pb/204Pb ⫽ 19.1 to 19.7 (Hoernle, unpublished data). In contrast, 187Os/188Os ratios are

U-series disequilibria in the Canary Islands

4159

Table 2. U-series concentrations and activity ratios measured by TIMS. Th U (238U)/ (230Th)/ (230Th)/ (230Th)/ (230Th)/ (234U)/ Ra (226Ra)/ (226Ra)/ Pa (231Pa)/ (231Pa)/ Age (yr) (ppm) (ppm) Th/U (232Th) (232Th) (232Th)o (238U) (238U)o (238U) (pg/g) (230Th) (230Th)o (pg/g) (235U) (235U)o Lanzarote EL1 EL10 EL11 EL5 EL3 EL13 EL8 EL15 EL17 EL18 EL21 EL20 EL23 Fuerteventura EF5 EF1 KFS240 EF10 EF7 EF9 EF12 EF6 Gran Canaria EGC7 EGC3 EGC6 EGC8 Tenerife ET2 ET7 ET6 ET8 ET1577 ET28 La Palma LP261925A LP71-7L LP124794 LP124793.1 ELP7 LP43e ELP1 WHJLP.92.11 Hierro EH8 EH11 EH4 EH17 TML

AD 1824 AD 1730 AD 1730 AD 1731 AD 1731 AD 1732 AD 1736 20,000 20,000 20,000 20,000 20,000 25,000?

6.93 6.03 6.04 3.32 2.16 2.34 1.96 5.03 4.49 4.91 5.72 5.40 4.46

1.74 1.59 1.62 0.84 0.56 0.60 0.52 1.40 1.19 1.30 1.47 1.28 1.11

4.0 3.8 3.7 3.9 3.8 3.9 3.8 3.6 3.8 3.8 3.9 4.2 4.0

0.763 0.800 0.813 0.772 0.788 0.778 0.799 0.847 0.801 0.806 0.778 0.721 0.756

0.981 1.034 1.044 0.990 0.939 0.980 0.948 1.269 1.167 1.143 1.101 0.922 0.868

0.981 1.034 1.044 0.990 0.939 0.980 0.948 1.354 1.241 1.202 1.166 0.963 0.897

1.286 1.293 1.284 1.284 1.191 1.259 1.186 1.499 1.458 1.409 1.415 1.278 1.148

1.286 1.293 1.284 1.284 1.191 1.259 1.186 1.600 1.550 1.492 1.499 1.335 1.186

1.004 1.001 1.007 1.001

1.306 0.763

1.700 1.084

1.755 1.137 1.991 1.095 0.985 1.892

1.995 1.897

0.458

1.220

1.247 0.515 1.863

1.868

1.005 0.996 0.997 1.005 1.008 0.992 0.995

0.292 0.731

1.389 1.015

1.436 0.294 1.740 0.471 1.026

1.744 1.040

0.636

1.005

5000? 5000? 5000? 5000? 5000? 5000? 5000? 5000?

4.10 2.54 6.55 3.87 3.69 2.50 2.46 2.11

1.00 0.68 1.66 0.88 0.92 0.61 0.46 0.41

4.1 3.8 4.0 4.4 4.0 4.1 5.3 5.1

0.733 0.806 0.768 0.690 0.759 0.741 0.567 0.593

0.926 0.953 0.923 0.872 0.877 0.840 0.761 0.715

0.932 0.959 0.930 0.881 0.882 0.846 0.771 0.721

1.260 1.181 1.202 1.265 1.156 1.135 1.342 1.205

1.272 1.190 1.211 1.277 1.163 1.142 1.359 1.215

1.001 1.003

0.198 0.165

0.936 0.970

3075 BP 3080 BP 3075 BP 10,000?

4.87 7.38 8.94 5.11

1.23 1.81 2.41 1.36

4.0 4.1 3.7 3.8

0.766 0.745 0.817 0.809

1.004 0.994 0.994 0.962

1.011 1.001 0.999 0.977

1.311 1.335 1.217 1.190

1.320 1.344 1.223 1.208

0.995 1.002 0.999 0.998

1.078

1.074

1.281 1.257 1.595

1.635

25,000? 4.96 25,000? 6.57 AD 1705 4.21 25,000? 5.19 AD 1909 4.87 800 BP 12.64

1.27 1.57 0.99 1.37 1.17 3.17

3.9 4.2 4.3 3.8 4.1 4.0

0.777 0.724 0.713 0.803 0.732 0.761

0.870 0.942 0.936 0.868 0.928 0.887

0.892 0.999 0.936 0.884 0.928 0.887

1.117 1.301 1.311 1.081 1.269 1.164

1.147 1.379 1.312 1.102 1.269 1.166

1.000 1.003 1.000 1.000 1.007 1.010

0.546 0.512 0.654 1.918

1.229 1.007 1.282 1.517

1.261 0.558 1.720

1.736

5000? 6.82 1.72 4.0 AD 1971 6.53 1.87 3.5 AD 1949 5.65 1.58 3.6 AD 1949 11.55 3.32 3.5 10,000? 7.31 2.03 3.6 AD 1646 9.08 2.40 3.8 AD 1480 7.78 2.20 3.5 2310 BP 19.80 6.20 3.2

0.766 0.870 0.849 0.873 0.841 0.802 0.858 0.950

1.033 1.107 1.106 1.112 1.065 1.053 1.141 1.177

1.046 1.104 1.105 1.112 1.086 1.053 1.141 1.183

1.348 1.268 1.302 1.274 1.266 1.313 1.330 1.239

1.365 1.268 1.302 1.274 1.292 1.313 1.330 1.245

0.999 1.006 0.999 1.003 0.997

1.266 1.014 2.079

1.552 1.439 1.433

1.558 1.060 1.734 1.448 0.878 1.695 1.442

1.734 1.696

1.615 1.379 3.146

1.486 1.377 1.197

1.565 1.330 1.692 1.471 1.238 1.718 1.535

1.697 1.720

AD 1793? 5000? 2900 BP 25,000? n⫽6

0.761 0.767 0.827 0.797 1.089

1.143 1.043 1.205 1.141 1.092

1.143 1.056 1.216 1.230

1.502 1.360 1.458 1.432 1.004

1.502 1.377 1.471 1.544

1.000 0.997 1.000 0.991 1.006

0.415

1.070

1.076 0.451 1.827

1.831

3.01 4.32 3.63 8.74 30.04

0.75 1.09 0.99 2.29 10.78

4.0 4.0 3.7 3.8

1.007 1.006 1.007 1.002 0.997

0.998 1.001

1.008

0.423 0.990 No corr.

1.293 1.731

Calibrated to 1.000

See text for analytical details. ( ) denotes activity. ␭238 ⫽ 1.551 ⫻ 10–10 y–1; ␭232 ⫽ 4.98 ⫻ 10–11 yr–1; ␭230 ⫽ 9.915 ⫻ 10– 6 yr–1; ␭226 ⫽ 4.326 ⫻ 10– 4 yr–1; ␭231 ⫽ 2.116 ⫻ 10–5 yr–1 ␭234 ⫽ 2.835 ⫻ 10– 6 yr–1. The ages were used in calculating initial (o) activity ratios. Question marks indicate age uncertainty when neither historical nor 14C dated; see text.

higher (⬎0.157) in the tholeiites and alkali basalts than in basanites (Widom et al., 1999). 4.2. U-Series Disequilibria Our results for U-series disequilibrium measurements are given in Table 2. In addition to providing the first 231Pa results for the Canaries, we provide the first comprehensive U-ThRa-Pa data set for Holocene volcanism throughout the Canary Islands. To assess reproducibility, we duplicated Th and U concentration analyses using separate digestions (Appendix B).

Th concentrations replicate within 1.4% with one exception (EL5, 2%), whereas the U analyses replicate within 1.1% with one exception (EG3, 1.9%). Errors based on the reproducibility of the analyses are ⫾0.7% for Th and ⫾0.5% (2␴) for U. Estimated error for Th/U based on eight replicates is better than ⫾0.7%. (230Th)/(232Th) replicated to within 1.3% 18 times with only 1 sample (ET8) outside the sum of the errors of the 2 analyses. (230Th)/(238U) replicated within error 8 times, (234U)/(238U) replicated to within 0.6% (n ⫽ 7 ) and (226Ra)/

4160

C. C. Lundstrom, K. Hoernle, and J. Gill

Fig. 2. MgO vs. SiO2 for all samples from this study. Two trends of silica enrichment are seen: one at lower MgO corresponds to the basanite to phonolite differentiation sequence observed in volcanic rocks from the western islands while the other corresponds to the sequence of near primitive basalts varying from basanite to tholeiite found on the eastern islands. The SiO2 increase from alkali basalt to phonolites is accompanied by decreasing Nb/U whereas the SiO2 increase defined by the basanite to tholeiite trend at relatively high MgO is also a trend of increasing Nb/U. The field at high SiO2 and low MgO corresponds to syenites from Tenerife (Wolff et al., 2001), all of which have Nb/U ⬍ 36.

(230Th) replicated within 0.9% (n ⫽ 5 ). The Pa concentration and (231Pa)/(235U) activity ratio for one sample replicated within 0.8%. Th concentrations range from 2 to 20 ppm in the whole sample set while U ranges from 0.5 to 6 ppm. Concentrations of Th and U steadily increase going from tholeiite to basanite to differentiated rocks and range from 4 to 9 and 1.0 to 2.4, respectively, in primitive basanites. Th/U of the whole sample set ranges from 3.2 to 5.4. (230Th)/(232Th)o varies widely from 0.88 to 1.35. The highest (230Th)/(232Th)o come from Volcan de Corona on Lanzarote for which the age given in Table 1 is an Ar-Ar isochron age (Carracedo, personal communication, 2002). On an equiline diagram, the entire data set produces a swath centered on (238U)/(232Th) of 0.78 (Fig. 4). All samples with Nb/U ⬎ 56 have relatively low (230Th)/(232Th)o and generally higher Th/U than other samples. Samples with Nb/U ⬍ 56 and low (230Th)/(232Th)o have the most uncertainty with regard to eruption age (lack of 226Ra excess). Later we will show that high Nb/U and high Th/U reflect incorporation of a lithospheric component (section 5.1). Ages older than assumed in Table 2 will result higher in (230Th)/(232Th)o and this may apply to samples ET2, ET8, EF9 and EH11. Removing high Nb/U and differentiated samples to examine solely the plume melting systematics produces a relatively tight cluster of data between (238U)/(232Th) of 0.72 to 0.87 with an average Th/U ⫽ 3.84 ⫾ 0.20. This narrow Th/U contrasts with the variable (230Th)/(232Th) and results in a large range of (230Th)/(238U)o (1.1–1.6). The highest (230Th)/(238U)o are found in the Volcan de Corona alkali basalts while most historical basanites have (230Th)/(232Th)0 of 0.94 to 1.10 and (230Th)/(238U)o of 1.3 to 1.5. For comparison, Hawaiian alkali basalts and basanites have (230Th)/(238U)o of only 1.15 to 1.30 (Sims et al., 1999). Thus, all samples show relative Th enrichment with an extremely large range in initial (230Th)/(238U). The (234U)/(238U) are all within error of 1.00 indicating no

Fig. 3. Selected incompatible element ratios as a function of U concentration for both Holocene and pre-Holocene samples (both primitive and differentiated) from the Canary Islands. The shaded portion of the Nb/U panel corresponds to Nb/U ⫽ 47 ⫾ 10 which are considered normal for MORB and OIB (Hofmann, 1980). The increase in these ratios with decrease in U content is interpreted to reflect addition of a lithospheric component (see text). The inset in the lower panel, a plot of Nb/Th vs. Th, shows the same systematics as plots with U as a denominator. This indicates that variations in (Nb, Zr, K)/U do not reflect mobilization of U during weathering. The complete overlap between Holocene and pre-Holocene data indicates the processes identified in recent samples have occurred continuously since the inception of the islands. The Fuerteventura basal complex rocks with high (Nb, K, Zr)/(Th, U) consist of carbonatites, ijolites, syenites, gabbros, and pyroxenites which have been fenitized and partially remelted to form migmatites (leucocratic veins) by later intrusives (Hoernle and Tilton, 1991). Pre-Holocene data from Pliocene samples from Hoernle et al. (1991a), Hoernle and Tilton (1991), Hoernle and Schmincke (1993a, 1993b), and Hoernle (unpublished data).

significant addition of secondary U to the samples which would affect (230Th)/(238U). (226Ra)/(230Th)o lie between 1.08 and 1.75 for samples of known eruption age, indicating significant parent-daughter dis-

U-series disequilibria in the Canary Islands

Fig. 4. Equiline diagram for all basalts in this study. All samples have excess 230Th over 238U. Samples with Nb/U ⬎ 56 have open symbols and three evolved (tephrite to tephriphonolite) samples with Nb/U ⬍ 36 have symbols containing crosses. We attribute the instances of coupled high Nb/U and high Th/U to addition of a lithospheric component. This component could be remelted wall rocks (orthoclase rich leucocratic veins) that have extremely high Th/U or could be Nb rich amphibole from the lithospheric mantle. Filtering both high and low Nb/U samples results in a restricted range in Th/U in melts (interpreted as the plume component) which ranges between ⬃20 and 60% excess 230Th.

equilibria despite the short half-life of 226Ra (1600 yr). This range is slightly larger than previous measurements of (226Ra)/ (230Th)o in Canary Island samples (Sigmarsson et al., 1998; Thomas et al., 1999) and other ocean islands. (226Ra)/(230Th)o is higher in tholeiites than basanites within the 1730 –1736 Lanzarote eruption but higher still in other basanites from Lanzarote and La Palma. Differentiated samples have (226Ra)/ (230Th)o as high as 1.7 (e.g., a tephriphonolite with Mg# ⫽ 44). The 226Ra deficits in two Fuerteventura alkali basalts may reflect either preferential retention of Ra during mantle melting or preferential leaching of Ra after eruption. (231Pa)/(235U)o shows a restricted range between 1.6 and 2.0 with the exception of two samples from the Volcan de Corona eruption on Lanzarote which have (231Pa)/(235U)o ⫽ 0.99 and 1.04. 231Pa excesses in basanites are generally higher in Lanzarote than the western islands and are slightly lower in tholeiites than basanites on Lanzarote. Considering primitive basalts only, two separate trends can be seen within the (230Th)/(238U) vs. SiO2 diagram (Fig. 5). One trend, defined by the Volcan de Corona lavas on Lanzarote, shows increasing 230Th excess as SiO2 increases from basanites to alkali basalts. We will refer to this later as a “plume” trend and attribute it to decompression melting within the plume. The other trend shows decreasing 230Th excess with increasing SiO2 and is best represented by the 1730 –1736 Timanfaya eruption on Lanzarote. Given the observed increase in Nb/U in these samples, we will refer to this later as a “lithospheric” trend and attribute it to lithospheric interaction and melting processes during melt ascent. 5. DISCUSSION

5.1. Variations in Nb/U: An Indicator of Lithospheric Interactions Variation in (Rb, Ba, K, Nb, Ta, Zr)/(La, U, Th) (section 4.1) provides critical insight into distinguishing the magma gener-

4161

Fig. 5. (230Th)/(238U)o vs. SiO2 for mafic volcanic rocks from the Canary Islands having Mg# ⬎ 60. Western island samples primarily are basanites having (230Th)/(238U)o of 1.2 to 1.5. Eastern island samples produce two distinct trends interpreted as two different effects of thinned lithosphere. Increasing (230Th)/(238U) with increase in SiO2 for the Volcan de Corona samples reflects a plume melting trend interpreted as an increasing length of a reactive porous flow melting column. A second trend of decreasing (230Th)/(238U) with increase in SiO2, defined by the Timanfaya eruption, reflects mixing between plume basanites and a lithospheric melt.

ation processes occurring in the Canary Islands. Hofmann et al. (1986) noted that Nb/U was relatively constant in OIB and MORB having a value of 47 ⫾ 10 and a more recent assessment by Sims and DePaolo (1997) found the mantle value to be 50 ⫾ 15. In contrast, Nb/U in Canary Island basalts has shown a consistent variation with U content from Miocene to present (Fig. 3) that is independent of U mobility. We identify and distinguish three separate processes based on Nb/U. First, because OIB and MORB lavas center on Nb/U of 47 ⫾ 10, samples within this window are interpreted to represent the plume melting process. Approximately half the primitive (Mg# ⬎ 60) samples lie within this range. Second, we interpret samples with Nb/U ⬎ 56 to reflect addition of a melt component from oceanic lithosphere. Nb/U extends to values as high as 111 as U and most other incompatible elements decrease (Fig. 3). Finally, Nb/U ⬍ 36 occurs primarily in differentiated samples from the western islands which we interpret to reflect crystallization of amphibole and other phases within the mantle lithosphere. Excellent linear correlation between Nb/U and Zr/U exists for the Canary Island sample set, consistent with binary mixing between the plume component and a high Nb/U, Zr/U component (Fig. 6). Two different oceanic lithospheric materials could serve as this component: (1) remelted intrusive rocks in the form of felsic (albitic to orthoclasic) veins, and (2) amphiboles and other minerals crystallized during differentiation of ascending plume-derived alkali basalts. In both cases, the high Nb/U signature of the lithosphere reflects the process of plume melts reacting with and changing the original lithospheric trace element signature, as has been observed in mantle xenoliths from the Canary Islands (e.g., Neumann, 1991; Kogarko et al., 1995; Whitehouse and Neumann, 1995; Wulff-Pedersen et al., 1996). The uplifted basal complex on Fuerteventura contains a wide range of plutonic rocks including alkalic gabbros, pyroxenites, syenites and carbonatites. These rocks have Sr-Nd-Pb isotopic compositions reflecting the Canary Island plume and have

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Fig. 6. Nb/U vs. Zr/U for all samples from this study (diamonds). The linearity of the basalt data is consistent with a mixing process controlling much of the variation. The inferred plume value is shown as an open square. Circles are amphibole compositions observed within a Tenerife basanite (Neumann et al., 1999) while squares are leucocratic vein compositions within the Fuerteventura basal complex (Hoernle, unpublished data). Both of these materials have extremely high Nb/U and Zr/U and could serve as possible lithospheric contaminants for elevating Nb/U in eastern island volcanic rocks.

extreme incompatible element contents and ratios: Nb/U ⬎ 8200, K/U ⫽ 825,000, Zr/U ⫽ 70,000 and Th/U ⬎ 100 (Hoernle and Tilton, 1991; Hoernle et al., 2002). Leucocratic (albite- and orthoclase-rich) veins have some of the most extreme ratios in the basal complex rocks, are generally rich in Si, Al, Na, Sr, Zr and Ba (plus K, Rb and Nb in the orthoclase veins), but poor in other major elements and trace elements. These veins are interpreted to have formed through the remelting of the older fenitized, gabbroic wall rocks in the contact metamorphic zones of large intrusions. Assimilation of the orthoclase-rich veins would lead to increased (Nb, K, Zr, Rb, Ba)/(Th, U) and Th/U ratios in the intruding magmas. A second source of high Nb/U that is unambiguously present is amphibole veins within peridotite. Enrichment in Rb, Ba, Nb, Ta, K relative to Th, U, and the light REE, and enrichment in Sr and Ti relative to the heavy REE in the high Nb/U basalts are consistent with assimilation of amphibole by ascending magmas (Ionov and Hofmann, 1995; LaTourrette et al., 1995). The presence of amphibole in this source would restrict it to the lithospheric mantle, since temperatures within the upwelling plume are too high for amphibole stability (Class and Goldstein, 1997). Vein amphiboles within composite xenoliths from southern Siberia and Mongolia (derived from continental lithosphere) have extremely high Nb and low U contents (Nb/U ⬎ 4400; Ionov and Hofmann, 1995). Amphibole-bearing veins in lithospheric xenoliths from the western Canary Islands reflect crystallization from plume-derived melts during ascent and have high Nb (up to 130 –150 ppm) but unknown U contents (Wulff-Pedersen et al., 1999). Amphiboles with reaction textures in Tenerife basanites average 330 ppm Nb and 0.3 ppm U, indicating that high Nb/U amphibole exists within Canary Island lithosphere (Neumann et al., 1999). Whatever the high Nb/U source is, it is not a long-lived feature as no consistent relationship exists between Nb/U and radiogenic isotope ratios (Fig. 7). This observation argues that

Fig. 7. Nb/U vs. 87Sr/86Sr (Hoernle, unpublished data) for all Canary Island basalts in this study. No clear correlation between Nb/U and radiogenic isotopes is observed indicating that increasing Nb/U does not reflect incorporation of isotopically distinct materials (such as heterogeneities within the plume or materials incorporated at the time of lithospheric formation at the ridge crest). Rather, the high Nb/U component must reflect relatively recent production in the lithosphere, interpreted as passage and reaction of previous Canary plume magmas.

high Nb/U does not reflect heterogeneity within the plume source or old lithospheric metasomatism (e.g., Halliday et al., 1995). Instead, the wide range in Nb/U occurs within the narrow variation in 87Sr/86Sr for the Canary Islands, consistent with the contaminant being produced by recent crystallization from ascending plume melts (Wulff-Pedersen et al., 1999). If the high Nb/U derives from the addition of high Nb/U minerals, then these minerals must completely dissolve in the melt without re-equilibration with the anhydrous residue (O’Hara et al., 2001). This is consistent with the lack of amphibole in eastern island xenoliths (Neumann et al., 1995). This situation is similar to observations from Gaussberg lamproites where the bulk trace element signature of a hydrous phase is imposed upon the melt rather than the signature of it being residual (Williams et al., 1992). Since high Nb/U samples are primarily alkali basalts and tholeiites from the eastern islands, the addition of the high Nb/U component is somehow linked to enrichment in silica and depletion in incompatible elements going from basanite to tholeiite. Our contention that lithospheric contamination results in relatively high Nb differs from previous work suggesting negative Nb anomalies in flood basalts reflected a subcontinental lithospheric component (e.g., Arndt et al., 1993, Turner and Hawkesworth, 1995). The latter typically have Nb/La ⬍ 1 whereas OIB, especially HIMU-rich ones, have Nb/La ⬎ 1. Indeed, Canary Island basanites have Nb/La ⬃ 1.3 and the samples with high Nb/U have still higher Nb/La (1.5–2). This may reflect a fundamental difference between oceanic and continental lithosphere. Finally, we interpret samples with Nb/U⬍36 to result from the crystallization of amphibole ⫾ associated mineral veins in oceanic lithosphere. The complementary nature of this process is seen by the linear extension beyond inferred plume values to low Nb/U and Zr/U in differentiated rocks (Fig. 6). Amphibole is a common phenocryst phase in tephritic through phonolitic rocks throughout the Canary Islands and is a required fractionating phase to derive these rocks from basanitic parents. Low Nb/U in differentiates of basanite in the western Canary Islands (i.e., in rocks with MgO ⬍ 4%) can be generated through crystal fractionation of alkalic magma during ascent through

U-series disequilibria in the Canary Islands

cold lithosphere. Our observation is reinforced by low Nb/U in “BDS syenites” from Tenerife which have an average Nb/U of ⬃21 (Wolff et al., 2000; 11 samples with Nb/U ⬍ 30 define the field shown in Fig. 2). Using a model similar to Klu¨ gel et al. (2000), amphibole fractionation can explain the observed decrease in Nb/U and increase in both Nb and U concentrations for the basanite (e.g., LP124794) to phonotephrite (e.g., WHJLP.92.11) sequence on La Palma if amphibole/meltDNb ⫽ 0.8. Although this DNb is higher than that measured by LaTourrette et al. (1995), it is within the range of those measured by Tiepolo et al. (2000). Furthermore, the presence of amphiboles having 330 ppm Nb in a Tenerife basanite (Neumann et al., 1999) indicates that amphibole/meltDNb can be ⬎ 1 because the maximum Nb content in most Canary Island magmas is ⬃250 ppm. The presence of amphibole vein xenoliths on the western Canary Islands (where differentiated rocks occur within the Holocene) is consistent with the interpretation that such veining is an intrinsic part of the process of alkali basalt differentiation (Neumann et al., 1995). Other minerals may also play a role in changing Nb/U. Phlogopite is also present in veins and its effect on Nb/U will be similar to amphibole’s (La Tourette, 1995). Nb is much more compatible in sphene or rutile, both common phenocryst phases in Canary evolved rocks and cumulus xenoliths and these phases may also be present in veins. So, although we refer to “amphibole veins” for simplicity, we recognize that other associated minerals may be as or more important in the differentiation and subsequent contamination process, and that these minerals form cumulates as well as veins. Although the actual mineral assemblage is important genetically, in this paper we merely infer that deviations from Nb/U ⫽ 47 ⫾ 10 result from an open system within the oceanic lithosphere. In summary, we hypothesize a complementary two-stage process to explain the observed variation in Nb/U. During the formation of a Canary Island, plume magmas ascend through thick, cold lithosphere crystallizing high Nb/U amphibole veins leading to the production of low Nb/U phonotephrites through phonolites (stage 1). In the Holocene, evidence for this process comes from the western (younger) islands in the form of amphibole veins in mantle xenoliths, abundant amphibole ⫾ sphene ⫾ rutile ⫾ phlogopite cumulate xenoliths, common amphibole phenocrysts in tephrites through phonolites and the presence of differentiated lavas having low Nb/U (for example sample WHJLP.92.11 from La Palma and ET28 from Tenerife). In stage 2, long-term heating and thinning of the lithosphere by ascending plume-derived melts leads to the complete melting of the leucocratic veins or the amphibole veins. This stage is currently observed on the eastern islands and results in the production of primitive tholeiitic magmas containing the high Nb/U trace element signature. The removal of amphibole during this stage is consistent with the absence of amphibolebearing mantle xenoliths on the easternmost islands. Both stages can be repeated during the evolution of a volcano resulting in enrichments and depletions of diverse age and character. For example, transitional tholeiitic basalts occur at the peaks and voluminous evolved volcanic rocks in the waning stages of both the Miocene and Pliocene Cycles on Gran Canaria (Hoernle and Schmincke, 1993a, 1993b). However, the overall thinning of lithosphere beneath the eastern islands is

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irreversible within the life of a volcano. We will next discuss these two stages in greater detail. 5.2. Stage 1: Differentiation Processes and Time Scales on the Western Islands Basanite to phonolite sequences are observed at ocean islands worldwide, yet we still need a clearer picture of where and how this differentiation process occurs. Typically the basanite-tephrite-phonotephrite-tephriphonolite-phonolite series is attributed to fractional crystallization of olivine, clinopyroxene, amphibole, Fe-Ti oxides, apatite and alkali feldspar. Fluid inclusion studies on differentiated rocks from La Palma indicate that differentiation occurred at mantle depths (Klu¨ gel et al., 2000) while petrological data from Tenerife samples indicate that differentiation takes place at 0.6 to 1.0 GPa (Ablay et al., 1998). Histograms of fluid inclusions within rocks throughout the Canary Islands indicate pressures of inferred magma stagnation of 0.55 to 1.0, 0.2 to 0.4, and ⬃0.1 GPa (Hansteen et al., 1998). Neumann et al. (1999) infer pressures equal to that near the oceanic crust–Canary volcanic rock interface. Our observations indicate that differentiation up to tephriphonolite occurs at a fast rate. Our samples include a basanite and tephrite pair from a compositionally zoned eruption in 1949 on La Palma previously studied by Sigmarsson et al. (1992). The changes in major and trace element composition of this sequence are consistent with crystallization of amphibole plus other phases. Whereas the 1949 basanite LP124794 has Nb/U of 39 and Ba/Th of 73, the 1949 tephrite LP124793.1 has lower Nb/U (31) and Ba/Th (65) while the nearby Nambroque phonotephrite WHJLP.92.11 has still lower Nb/U (23) and Ba/Th (52). Hawkesworth et al. (2000) also interpreted the depth and time scales of alkalic differentiation on a western Canary Island (Tenerife) using techniques similar to this study. However, their conclusions differ significantly from ours by inferring a process of differentiation in which basaltic melts pond at the base of the crust for up to 230,000 yr. This time scale reflected the decrease in (230Th)/(238U) from 1.5 to 1.1 with increasing incompatible element (Zr) concentration as basanites evolve to phonotephrites by 50% fractional crystallization. Our combined 230 Th-226Ra-231Pa data are inconsistent with a long time scale for phonotephrite formation. Although (230Th)/(238U) of our samples decreases as Zr content increases similar to Hawkesworth et al. (2000), (226Ra)/(230Th) is ⬎ 1 for all samples, inconsistent with a long time scale for differentiation (Fig. 8). Indeed, tephriphonolite ET28 from Tenerife (at least 600 yr old) has (226Ra)/(230Th) of 1.52, resulting in one of the highest (226Ra)/(230Th)o. Phonotephrite WHJLP92.11 (⬃2310 yr old; Carracedo, unpublished data) has a (226Ra)/(230Th)o of 1.53, similar to basanites from the 1646 and 1971 eruptions on La Palma. Excess 226Ra argues that no magma experienced a differentiation time ⬎ 8000 yr. Furthermore, the relative constancy of (226Ra)/(230Th)o as Zr concentration doubles in La Palma samples provides robust evidence for fast fractionation all the way to tephriphonolite, involving solidification of as much as 75% of the melt within a few centuries. Thus, although (230Th)/(238U) alone could be interpreted to mean differentiation lasted 230,000 yr (Hawkesworth et al.,

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Fig. 8. (226Ra)/(230Th)o and (230Th)/(238U)o as a function of Zr content for samples from the western Canary Islands. The increase in Zr corresponds to decreasing Mg# with differentiation. Gray field represents data from Hawkesworth et al. (2000) who interpreted the decreasing (230Th)/(238U)o to reflect differentiation from basanite to trachyte taking 230,000 yr. Although our (230Th)/(238U)o data broadly follow those of Hawkesworth et al. (2000) and therefore could be interpreted as the time of differentiation, (226Ra)/(230Th)o will decay back to 1.0 over a time scale of ⬃8000 yr. The presence of 226Ra excess in every sample clearly contradicts this interpretation implying a much shorter time scale for differentiation (hundreds of years). Error bars are similar to symbol size.

2000), this hypothesis is inconsistent with the other parentdaughter pairs from similar differentiation suites in the Canary Islands. It is also inconsistent with petrologic work showing that the differentiation process beneath Tenerife involves an open system of repeated mixing between older fractionated magmas and younger mafic magma (Neumann et al., 1999). A more consistent explanation is that the process of basanite differentiating to phonolite occurs by amphibole (⫾ other minerals) crystallization during ascent of plume-derived basalts through the lithosphere. In this interpretation, the time scale of differentiation reflects the transport time of an individual magma batch through the lithosphere. 226Ra excesses constrain the time of final transport from the location of last crystal-melt equilibrium to the surface to be within a few hundred years. This is consistent with differentiation from basanite to phonotephrite within 13 yr before the 1949 La Palma eruption based on seismic activity preceding the eruption (Klu¨ gel et al., 2000). 5.3. Stage 2: The Generation of Primitive Basanites Through Tholeiites on the Eastern Islands The amount of 230Th excess as a function of SiO2 varies along two distinct trends distinguished by differences in Nb/U

(Fig. 5). Both trends emanate from the locus of primitive basanites with SiO2 of ⬃42 wt.% and (230Th)/(238U) of ⬃1.2 to 1.4 which occur on both eastern and western islands. However, eastern island samples define the extension of each trend to higher SiO2. In the first trend, exemplified by the 1730 –1736 Timanfaya eruption but encompassing many of the samples, 230 Th excesses decrease from basanite to tholeiite coincident with a progressive increase in Nb/U (as well as [Rb, Ba, K, Nb, Ta, Zr]/[U, Th, LREE]). A second trend, defined by samples from the Volcan de Corona eruption on Lanzarote, shows 230Th excess increasing as SiO2 increases to alkali basalts. The latter samples all have Nb/U of 47 ⫾ 10 and are therefore interpreted to reflect melting solely within the plume. We compare these two trends below after discussing previous models of magma genesis in the Canary Islands. The diversity of primitive magmas having broadly similar Sr, Nd, and Pb isotopic compositions would appear to make the Canary Islands an excellent setting to study variations in degree of melting in a slowly upwelling plume. Indeed, basanites through alkali basalts through tholeiites, all containing mantle xenoliths, can be found within one eruption sequence, the 1730 –1736 Timanfaya eruption on Lanzarote. The presence of xenoliths and high MgO (9.3–14.3 wt.%), Ni (269 – 440 ppm), Cr (368 –580 ppm) and Mg# values (66 –73) testify to the near-primary nature of these magmas and to their very short crustal residence times. We use this eruption as a template for understanding the systematics of eastern island magmatism The Timanfaya eruption has been the focus of several previous studies, providing three distinct interpretations. Although all studies concur that two-component mixing explains the geochemical trends from basanite to tholeiite, the explanation of how the endmember melts are created varies widely. For instance, Sigmarsson et al. (1998) argued that basanites reflected relatively large degrees of batch melting (10%) of garnet pyroxenite veins whereas tholeiites were produced by smaller degree melting (1%) of garnet lherzolite. Thomas et al. (1999) concluded that basanites through tholeiites reflected increasing degrees of melting of homogeneous plume peridotite source as pressure decreased slightly (final depths of melting of 60 –70 km). Reiners (2002) proposed that the basanites reflected low-degree (0.1–5%) melting of plume (or possibly lithospheric) lherzolite while tholeiites reflected high-degree (⬃20 – 60%) melting of a plume pyroxenite source. Indeed, the wide range in proposed hypotheses is evidenced by the exact opposite interpretations of the Sigmarsson et al. (1998) and Reiners (2002) models. These models explain some of the features of the Timanfaya eruption but cannot adequately explain the full range of geochemistry. We emphasize that none of the models can explain the atypically high (Nb, K, Ba, Rb)/(Th, U, LREE) ratios observed in more silica rich basalts from the eastern islands unless these ratios reflect plume heterogeneity; the lack of any clear correlation between these ratios and long-lived isotope ratios (e.g., Fig. 7) argues against this. Further, the high and near constant FeO contents of Timanfaya basalts over a range in SiO2 (43–51 wt.%) and MgO are inconsistent with experimentally predicted peridotite melting behavior (Fig. 9). For instance, generation of the Timanfaya tholeiites by peridotite partial melting appears to require pressures ⬍ 1 GPa and a source more fertile than HK66 (Mg# ⬍ 85) (Hirose and

U-series disequilibria in the Canary Islands

Fig. 9. FeO vs. SiO2 for basalts from the Timanfaya eruption (Mg# ⬎ 60; squares). Circled fields represent the experimental data of Hirose and Kushiro (1993) at a single pressure (in GPa) for two different peridotite starting materials. Underlined numbers denote pressure of experiments on HK-66 (Mg# ⫽ 86) while those without underlining reflect pressure of experiments on KLB-1 (Mg# ⫽ 89). Each field represents two to three experiments at different temperature; therefore variations in degree of melting at a single pressure are contained within the field shown for that pressure. The data from the Timanfaya eruption or the Canary Islands in general are not consistent with either variations in pressure, variations in peridotite source composition or variations in degree of melting. Although experiments predict a strong polybaric correlation between FeO and SiO2, the Timanfaya samples have high and near constant FeO despite SiO2 varying from 43 to 51 wt.%. Mixing between plume-derived basanites and either of two possible melts from the lithosphere might explain the observed systematics. One possible melt is a high silica melt created by alkali diffusive infiltrationthe arrow points to a high silica melt of 63 wt.% SiO2 and 6.7 wt.% FeO (Table 3) produced by diffusive infiltration into harzburgite (Lundstrom, in press). A second possible explanation is direct reaction between the plume-derived basanite and lithospheric mantle (Table 3).

Kushiro, 1993). Melts of mixtures of 50 to 67% peridotite (KLB-1) and 33 to 50% MORB have FeO contents generally lower than observed and very different CaO, Na2O and Al2O3 (Kogiso et al., 1998). These discrepancies argue against either decompression melting with small variation in depth (Thomas et al., 1999) or melting of lherzolite at plume depths to produce the tholeiites (Sigmarsson et al., 1998). Even hydrous melting of lherzolite, for which there is little evidence for in Timanfaya, cannot produce the FeO observed in the tholeiites at the observed MgO (Hirose and Kawamoto, 1995). Reiners’s (2002) model of tholeiites reflecting high degree melting of a pyroxenite source has merit given recent results of eclogite melting experiments (Pertermann and Hirschmann, in press). These show that 60% melting of pyroxenite at 3 GPa can produce melts with high SiO2 and Al2O3 but low CaO and MgO, similar to the contents observed in the Timanfaya tholeiites (Reiners, 2002). Trace element models of mixing melts from pyroxenite and lherzolite can reproduce the observed systematics although these depend on assumptions about eclogite source trace element composition which is largely unconstrained. However, in the context of the Canary Islands as a whole, the Reiners (2002) model fails to explain the exclusive presence of basanites and lack of tholeiites on the younger, western islands. If the plume source reflects a concentration of eclogite, then tholeiitic magmas should be found throughout the islands and particularly in western islands where the plume is

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likely centered. Further, one of the main arguments used by Reiners (2002) to support a pyroxenite source for the tholeiite vs. a lherzolite source for the basanite in the Timanfaya eruption is 187Os/188Os varying from 0.129 to 0.157 and positively correlating with SiO2. However, none of the measured 187Os/ 188 O are thought to reflect the Canary Island plume (Widom et al., 1999). Instead, the ratios of the basanite and alkali basalt have been reduced by assimilation of xenocrystic olivine from the lithosphere whereas that of the tholeiite (which has a low Os concentration) has been raised by assimilation of sediment. The plume is thought to have an intermediate 187Os/188Os indicating the presence of 25 to 35% recycled oceanic crust in the source which elevates ratios above those in depleted mantle. Between this paper, Sigmarsson et al. (1998), and Thomas et al. (1999), there are 40 U-series analyses of lavas from this eruption, making it the best studied eruption for U-series disequilibria globally. Decreasing 230Th and 231Pa excess and increasing 226Ra excess from early basanite to late tholeiite characterize this eruption. Like other geochemical data, the U-series disequilibria are fully consistent with mixing playing a major role in controlling the disequilibria variation in the Timanfaya eruption. Rare earth element ratios and U-series disequilibria data distinguish melting of sources with garnet from those without garnet, thus placing constraints on the melting depths of the two end members. High (Sm/Yb)N and large 230Th excesses in basanites are consistent with derivation from a garnet-bearing source (Beattie, 1993). The age-corrected 230Th excesses of 60% in the Volcan de Corona eruption are some of the highest yet measured in any volcanic rock and are inconsistent with the slight excesses explicable solely by clinopyroxene-melt partitioning (Blundy et al., 1998; Wood et al., 1999; Landwehr et al., 2001). The mineralogical constraint of garnet is thus consistent with the expectation that melting to produce basanites occurs deeper than the lithospheric lid in the Canary Islands which on the western islands may be deeper than 100 km. For primitive samples, (Sm/Yb)N and CaO/Al2O3, as well as (230Th)/(238U), correlate inversely with SiO2 indicating a reduced garnet signature in the tholeiites. Although the tholeiites still have a garnet signature, the ubiquitous evidence for mixing argues that the role of garnet in the source of the tholeiites may be an artifact of the mixing. Finally, we note that basanites from both the western and eastern islands show wide variation in (226Ra)/(230Th)o (e.g., Lanzarote samples 1730 Timanfaya basanite vs. 1824 basanite: see Fig. 8). One possible explanation for this variation is that plume basanites pond for varying times as they pass through the lithosphere leading to variable transport times (Thomas et al., 1999) although the low viscosities of these magmas may argue against this. Alternatively, this variation may just reflect the inherent sensitivity of 226Ra excess generation to small variations in melt porosity as shown by ingrowth melting models (Spiegelman, 2000). Clearly, a better understanding of the process creating (226Ra)/(230Th) variation in the plume basanites is needed for constraining the times scales of basanite formation and ascent. The thinning of lithosphere beneath the eastern islands (section 2) has two effects important to the SiO2-230Th systematics (Fig. 5). First, thinner lithosphere allows the plume to upwell and melt to shallower depths resulting in higher silica content melts (the basanite-alkali basalt “plume

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trend”). Second, because alkalis have a greater effect on melt SiO2 at lower pressure (Hirschmann et al., 1998), melting of the lithosphere by interaction and reaction with basanites is more likely to occur (the “lithospheric trend”). We next discuss these two melting regimes. 5.3.1. The plume trend As discussed above, a variety of different models for creating the major element compositions of Canary Island magmas have been proposed. The origin of the plume basanites remains open to debate given the diversity of melts produced in experiments. For instance, experimental melting of pyroxenite can produce both silica-saturated compositions (Pertermann and Hirschmann, 1999) and silica-undersaturated compositions (Kogiso et al., 1998; Hirschmann, 2003), supporting the antithetical models of Sigmarsson et al. (1998) and Reiners (2002). Indeed the major element composition of a magma ultimately erupted may depend as much on the physical ability of a magma to ascend without reaction with surrounding peridotite as it does on the composition of its original mafic source (Lundstrom et al., 2000; Hirschmann, 2003). Melting experiments on mixtures of peridotite and eclogite show that that melts from eclogite quickly react with peridotite (Yaxley and Green, 1998), leading to a hybridized peridotite with phase compositions approaching those of peridotite. Based on Os isotopes for Canary Island basanites which argue for 30% pyroxenite in the plume source (Widom et al., 1999), we interpret basanites from the Canary Islands to most likely reflect melting of mafic pods within the plume source, in agreement with Sigmarsson et al. (1998) and Hoernle (1998), with peridotite hybridization possibly also being important. The combined U-Th-Ra-Pa data allow assessment of different models for the creation of U-series disequilibria during decompression melting within a plume. A distinctive characteristic of the Canary Islands plume samples is the large variation in (230Th)/(232Th) with small variation in (238U)/(232Th). Both simple mass balance melting models (batch or fractional melting) and more complex two-phase flow (ingrowth) models have been proposed as explanations for the generation of Useries disequilibria. Elliott (1997) reviewed the production of U-Th disequilibria in OIB and MORB and suggested that (230Th)/(238U) in OIB were consistent with production of 230Th excess by both degree of melting and ingrowth during the melting process. The original explanation for the ubiquitous 230Th excess in young volcanic samples was preferential extraction of Th relative to U from the mantle leaving behind a complementary low Th/U residue. Thus, if Th were more incompatible than U, then partial melts will have Th/U greater than the original source. In this model, the (230Th)/(232Th) of the source and melt are identical (e.g., Gill and Condomines, 1992). However, primitive basalts with 36 ⱕ Nb/U ⬍ 56 have twice as much variation in (230Th)/(232Th) as (238U)/(232Th) (Fig. 4). This observation alone is inconsistent with degree of melting variations since either batch or fractional melting would predict a horizontal, not vertical, array given a homogenous source. In addition, (230Th)/(238U) for samples with Nb/U ⬍ 56 correlates inversely with incompatible element concentrations (Fig. 10), opposite to the effects predicted by variable degrees of melting.

Given the incompatibility of the trace elements shown, the ⬃50% decrease in their concentration corresponds to doubling the degree of melting in going from the highest concentration basanites of Gran Canaria and La Palma to the Hierro basanites and the Volcan de Corona alkali basalts. A second explanation might be that the large range in (230Th)/(238U)o could reflect similar and large initial disequilibria in melts formed by a uniform process (uniform source and degree of melting, or uniform in-growth melting) followed by variable transport and/or differentiation times which range from zero in melts with the largest excess 230Th to 230,000 yr in those with none (Hawkesworth et al., 2000). As shown in section 5.2, some of our samples with the lowest (230Th)/(232Th)o are minimally differentiated and have significant 226Ra-230Th disequilibria, precluding long differentiation or transport times. Our preferred interpretation is that the range in (230Th)/ 232 ( Th)o represents variations in the amount of ingrowth of 230 Th within the melt column, reflecting the process of twophase flow during melting (McKenzie, 1985; Spiegelman and Elliott, 1993). In this explanation, melts can have higher (230Th)/(232Th) than their source because 230Th is produced by decay of its U parent which is preferentially retarded within the melting column. Once the degree of melting exceeds that needed to fractionate Th from U (⬎2% melting if DU ⬍ 0.01), all 230Th-238U disequilibria reflect ingrowth. Therefore, large variations in (230Th)/(232Th)o with little variation in Th/U are consistent with an ingrowth model if the degree of melting is ⬎ 2%. A longer equilibrium porous flow melting column (Spiegelman and Elliott, 1993) will result in higher total degrees of melting, lower incompatible element concentrations and can produce higher 230Th excess (Lundstrom et al., 1994). However, our model results show that reasonable variations in melting column length keeping all other parameters equal cannot account for the large range in (230Th)/(238U)o. If the melting process in the plume reflects reactions between melts of mafic lithologies and surrounding peridotite (hybridization), then the U-Th partition coefficients could change from those relevant to pyroxenite to those expected for peridotite melting (Stracke et al., 1999). If so, the melts from longer melt columns would produce larger 230Th excesses as well as lower incompatible element concentrations (Fig. 10). The higher SiO2 contents and lower Sm/Yb for the Volcan de Corona alkali basalts are also consistent with shallower depths of melt extraction. This model of variable depths of final melting is consistent with general petrologic observations within the Canary Islands. Basanites, derived from depths of melting equal to or greater than the thickness of 150 to 180 Ma oceanic lithosphere, are generated under all islands. However, alkali basalts and transitional tholeiites such as those from Volcan de Corona (eastern islands) form only in locations where the lithosphere has been thinned significantly allowing the plume to reach more shallow depths beneath eastern islands. In a variable length melting column model, the Hierro basanites reflect a longer melt column and greater percent melting than other basanites inconsistent with reflecting lithospheric thinning. Speculatively, the Hierro source may have had a slightly higher volatile content and so began to melt earlier and deeper than the other basanites. Alternatively, (230Th)/(238U)o across the archipelego could reflect variations in solid upwelling rate. Considering the is-

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Fig. 10. Incompatible trace element concentrations as a function of (230Th)/(238U)o for basalts with Nb/U ⬍ 56. Samples denoted by open circles have 52 ⬍ Nb/U ⬍ 56 and are interpreted to have had both trace element concentrations and (230Th)/(238U)o lowered by addition of lithospheric component. The sample with the lowest (230Th)/(238U)o, ET8, has little age constraint and could be old. The remaining samples, interpreted to reflect the plume melting process, show an inverse correlation between incompatible element concentration and 230Th excess. Samples to the right of the dashed line are from Hierro and Volcan de Corona, Lanzarote. An explanation for these trends is an increasing melting column length and higher degree of melting causing dilution of incompatible trace element concentrations and increasing 230Th excess by ingrowth melting. The two open squares reflect reactive porous flow melting models (Spiegelman, 2000) where melting starts at 4.0 GPa and ends at either 3.0 (higher concentration square in each panel, F ⫽ 0.03) or 2.0 GPa (F ⫽ 0.06) These models, using otherwise identical parameters (␾max ⫽ 0.2%, bulk DU ⫽ 0.006 and DTh⫽0.004), fail to match the observed range in (230Th)/(238U)o. However, if melting involves reaction and peridotite hybridization such that the partition coefficients increase (as the matrix becomes more peridotitic, Stracke et al., 1999: bulk DU ⫽ 0.01 and DTh ⫽ 0.006), larger excesses can be created, consistent with the observed trend (open square with slash).

lands with historic activity, higher 230Th excesses on the most exterior islands surround a trough of low excess at La Palma and Tenerife. Although each island shows considerable variation in (230Th)/(238U), the overall pattern is consistent with higher solid upwelling rates beneath La Palma and Tenerife, the inferred center of the hotspot, and lower upwelling rates on the plume edges. This plume center is consistent with La Palma and Tenerife lavas having the strongest Canary Island plume isotopic signature. Such a model is similar to that of Bourdon et al. (1998) who used variations in solid upwelling rate to explain a similar range in (230Th)/(238U)o. A difficulty with this model is explaining why slower upwelling rates beneath Hierro and Lanzarote result in lower incompatible element concentrations (Fig. 10). Indeed, both upwelling rate variations and melting column length variations may well occur within the Canary Islands, possibly leading to the less systematic behavior within the U-series data. 5.3.2. The lithospheric trend: The generation of tholeiitic melts Whereas the diversity of basanites erupted on the western islands indicates variations within the plume melting column,

the trend toward silica enrichment and low 230Th excess on the eastern islands reflects addition of lithospheric components (Fig. 5). We show that the trends of major, trace, and U-series elements in eastern island volcanics are consistent with mixing between plume-derived basanites and either of two possible melts formed in the lithosphere as a result of reaction between basanites and the mantle lithosphere. One possibility is that plume basanites stagnate and directly react with mantle lithosphere, primarily through incongruently dissolving orthopyroxene from the lithosphere. Piston cylinder experiments at 0.9 GPa and 1300°C using a mixture of basanite EL10 and spinel lherzolite KLB-1 produce melts that resemble tholeiite EL8 in SiO2 but have lower FeO (Table 3) (Lundstrom, submitted). The melt composition produced in these experiments is independent of the basanite to peridotite ratio for ratios ranging from 1:2 to 1:10. However, higher ratios of basanite to peridotite, which result in more iron rich bulk compositions, will produce higher melt FeO contents (analogous to comparison of melts of KG1 and KG2 in Kogiso et al., 1998). The EL10-KLB-1 experiments show that reaction and approach to melt-mineral equilibrium occurs rapidly (Lund-

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C. C. Lundstrom, K. Hoernle, and J. Gill Table 3. Observed and calculated Timanfaya melt compositions.

SiO2 TiO2 Al2O3 FeOt MnO MgO CaO Na2O K2O P2O5 a

EL-10

High silica melt from Harz.DIA experimenta

Calculated 60:40 mix EL-10/high silica melt

EL-8

Melt from direct reaction of EL 10-KLB-1a

42.86 2.98 11.60 11.82 0.20 14.24 10.07 3.72 1.51 0.98

62.96 0.71 12.90 6.72 0.11 6.40 5.21 5.20 0.97 0.14

50.90 2.07 12.12 9.78 0.16 11.11 8.12 4.31 1.29 0.65

50.74 2.44 13.47 10.49 0.15 9.47 9.19 3.08 0.63 0.34

50.55 2.13 16.09 7.32 0.13 6.42 11.62 2.89 1.01 0.56

Lundstrom (in press).

strom, in press), consistent with previous studies documenting fast dissolution of orthopyroxene in alkalic melts (Brearly and Scarfe, 1986; Shaw, 1999). Thus if basanites stagnate, as possibly implied by the variability in (226Ra)/(230Th), then direct reaction between basanite and the lithosphere is likely to occur resulting in tholeiitic melts with relatively high silica and iron. A second hypothesis is that the tholeiites reflect mixing between plume basanites and a high silica melt created in the lithosphere by the diffusive infiltration of alkalis (DIA). In this process, alkalis from ascending plume basanites diffuse into the surrounding lithosphere at 0 to 1 GPa causing orthopyroxene to melt incongruently (Lundstrom, 2000; Lundstrom, in press). Melts produced in DIA experiments resemble high silica glasses observed ubiquitously in Canary Island xenoliths and inferred to have a lithospheric mantle origin (e.g., Neumann et al., 1995; Wulff-Pedersen et al., 1999). Although the origin of high silica glasses in xenoliths remains to be established (see list of suggested origins in Draper and Green, 1997), we assume here that the DIA process can create the observed xenolith glasses and lithospheric melts of similar composition. This leads to a self-consistent model whereby plume basanites ascend in a melt conduit and lose alkalis by diffusion into the surrounding lithosphere creating high silica melt. The high silica melt then mixes with basanite producing tholeiites observed at the surface. Laboratory experiments juxtaposing basanite EL-10 with partially molten KLB-1 show that alkalis rapidly diffuse from the basanite into the lherzolite producing high silica melts within lherzolite (Lundstrom, 2000), consistent with the known effect of alkalis on the orthopyroxene-olivine phase boundary (Kushiro, 1975). DIA experiments with harzburgite produce melts with 63 wt.% SiO2 and relatively high FeO contents (Table 3) (Lundstrom, in press). Mixing of such melts with El10 in a 40:60 ratio produces a melt similar in composition to observed tholeiites including reproducing the FeO-SiO2 trend of the Timanfaya data (Fig. 9). The differences between the calculated mixture and EL8 could reflect (1) differences in the actual composition of the harzburgite beneath Lanzarote and that used in the experiments and (2) further mineral melt reactions and crystallization of melts during lithospheric ascent. We can evaluate the mixing hypothesis further by plotting trace element data for the Timanfaya sequence (1730 –1736 eruption) with ion probe analyses of high silica glasses found in

the interiors of xenoliths from La Palma, possible representatives of Canary Island lithospheric melts (Wulff-Pedersen et al., 1999). For diffusively immobile trace elements like Ti, Y, Sm, Hf, Yb, V, Zr, Cr, and Ni, good linear relations (R2 ⬎ 0.96) are observed between the Timanfaya lava suite and the xenolith glasses, consistent with our hypothesis (Fig. 11). An important observation is that the trends for many elements, including some that we use to identify lithospheric contamination, point to trace element concentrations of zero at 65 to 70 wt.% SiO2. These linear relationships occur in elements ranging from compatible (Cr, Ni) to incompatible (Zr, Hf). Since there is no apriori reason why the basanite-alkali basalt-tholeiite samples should consistently point toward an end member with zero concentration, these trends may identify a fundamental aspect of the basanite to alkali basalt to tholeiite relationship. In this regard, either a high silica glass created by DIA or a melt reflecting incongruent dissolution of orthopyroxene in basanite will lead to a high silica end member melt with low contents of both incompatible and compatible trace elements. Notably, when elements that are diffusively mobile are considered, linearity is lost; the xenolith glasses have alkali concentrations that are far too high to represent a mixing end member. This difference in elemental behavior is consistent with what might be expected of the DIA process. Alkali and alkaline earth elements, which have high diffusivities and low activity coefficients in silica-rich melts, will diffuse into the xenolith interiors while slower diffusing elements will not. In the case of Figure 11, the xenolith glass alkali concentrations reflect the ability of alkalis to rapidly diffuse from the host basalt to xenolith interior, whereas the process of melting lithosphere over longer diffusive distances will not result in alkali enrichments other than Na. This mixing hypothesis is also consistent with the inferred changes in source mineralogy for the two end member melts. The basanite end member has a REE pattern consistent with residual garnet whereas Sm and Yb concentrations in high silica glasses indicate a flat REE pattern, consistent with spinel lherzolite partitioning (Wulff-Pedersen et al., 1999). Mixing of lithospheric melts having (Sm/Yb)N ⫽ 1 with basanites having (Sm/Yb)N ⫽ 7 results in (Sm/Yb)N similar to those observed in the tholeiites, 4. This same mixing calculation explains the variations in (230Th)/(238U). At low pressure in the presence of clinopyroxene, the (230Th)/(238U) in a lithospheric melt should

U-series disequilibria in the Canary Islands

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Fig. 11. Variations in trace element concentrations as a function of SiO2 for samples from the Timanfaya eruption along with high silica glasses found in mantle xenoliths from La Palma (open symbols: interior glass samples with SiO2 ⬎ 65 wt.% from Wulff-Pedersen et al., 1999). Concentrations for Hf and Ni are assumed to be zero in the high silica glass endmember. (A, B) Diffusively immobile trace elements form highly correlated trends for the basanite-alkali basalt-tholeiite sequence combined with the high silica glasses. Based on this, tholeiites are consistent with mixing between plume-derived basanites and a high silica melt from the lithosphere. (C, D) In contrast, diffusively mobile elements like alkali and alkaline earth elements do not form the same linear relationships with the high silica glasses. Instead, high silica glasses have higher concentrations of these elements than tholeiites and sometimes higher than basanites. The difference in behavior between elements in A and B relative to C and D is attributed to the relative diffusion rates of each element through silicate melts with diffusively mobile elements preferentially enriched in the xenolith glasses (see text). This explanation is consistent with measured basanite melt-tholeiite melt diffusion coefficients which show that alkalis and alkaline earth elements diffuse an order of magnitude faster than high field strength elements (Lundstrom, in press).

be ⬃1.0 (LaTourrette and Burnett, 1992; Beattie, 1993; Lundstrom et al., 1994; Landewehr, 2000). The (230Th)/(238U) of 1.19 observed in the Timanfaya tholeiite is consistent with 60:40 mixing between the observed basanites (1.3) and a lithospheric endmember which has (230Th)/(238U) ⫽ 1.0 (Fig. 12). Thus, although the tholeiites have “garnet” signatures of 230Th excess and (Sm/Yb)N ⬎ 1, this signature simply reflects the garnet signature of the basanite which makes up 60% of the mixture. The same basanite-high silica glass mixing relationship suggests a (226Ra)/(230Th) of ⬎ 2 for the lithospheric end member (Fig. 12). The lithospheric signature of higher (K, Nb, Ta, Zr, Rb, Ba)/(LREE, U, Th) is not explained by mixing with a highsilica melt because most of these elements are absent from that melt. Rather, for these elements, the primary effect of mixing is dilution, making contamination by amphibole veins or migmatites more effective (section 5.1). The only connection between the DIA process and the contamination is that both processes reflect interactions between ascending plume melts and the lithosphere. The occurrence of DIA in the eastern Canary Islands should be expected for two reasons. The long time scale over which plume magmas have passed through eastern island lithosphere will raise the temperature of the peridotite wall rock surrounding melt conduits, allowing alkalis to flux into a considerable volume of lithosphere. Since the lithosphere beneath the eastern islands has experienced one of the longest periods of repeated alkali volcanism anywhere in the oceans (perhaps 60 – 80 Ma;

Le Bas et al., 1986; Balogh et al., 1999), it is a prime target for being affected by DIA. Second, the eastern islands appear to have experienced significant lithospheric thinning (section 2). The DIA process will become most prominent at ⬍ 1 GPa pressures since melt alkali contents have a much greater influence on melt silica contents at low pressure (Hirschmann et al., 1998). Could the DIA process result in the volume of tholeiitic melt observed in the 1730 –1736 eruption (at least 10x more tholeiite than basanite)? If basanites have fluxed through the lithosphere beneath Lanzarote for the last 25 and possibly 60 to 80 Ma, then significant amounts of alkalis could have diffused into the lithosphere making it alkali rich and significantly decreasing its melting point. As a simple example, assume an ascending plume-derived melt forms a dunite melt conduit through the lithosphere that is 0.5 km in radius and 50 km in length and has a melt porosity of 5% during basanite ascent. After a plumederived basanite ascends through it, the melt conduit is surrounded by 150 m of diffusively fluxed peridotite that undergoes 5% melting. The mix of these two melts (60% basanite to 40% high silica melt; Table 3) results in 3.3 km3 of tholeiite, similar to the observed Timanfaya eruption volume of 3 to 5 km3 (Carracedo et al., 1990). Although this scenario is oversimplified as the basanites more likely ascend through multiple conduits, it indicates that a relatively small diffusively fluxed rind could account for the mixing hypothesis. It remains difficult to understand how the viscous, high silica melt in the area surrounding a conduit can be extracted and mixed with the

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5.4. Comparison of U-Series Disequilibria in Oceanic Mantle Settings High precision mass spectrometric data for (226Ra)/(230Th), ( Pa)/(235U), and (230Th)/(238U) now exist for a number of MORB and OIB suites. Because the disequilibria of each parent-daughter pair is primarily influenced by two unique parameters, the decay constant of the daughter nuclide and Dparent/Ddaughter, each parent-daughter pair provides different information about the melting process. Numerical models of two phase flow show that 231Pa and 226Ra are more sensitive to variations in porosity than is (230Th)/(238U) (Spiegelman and Elliott, 1993). The short half-life of 226Ra (1600 yr) means that excesses created very deep in the melting column are unlikely to be preserved whereas the half-lives of 231Pa and 230Th are long enough to reflect the entire melting column; thus these two nuclides are sensitive to increases in melting column length. However, if garnet controls U-Th partitioning (Beattie, 1993), 230 Th excesses will not reflect the entire melt column because garnet will not be stable at depths shallower than 90 to 50 km, depending on composition. Thus, comparison of disequilibria variations within OIB and differences in disequilibria between OIB and MORB provides key insight into the process of melting oceanic mantle. Canary Island mafic volcanic rocks have (230Th)/(238U) and 231 ( Pa)/(235U) similar to disequilibria observed in the Comores (1.3–1.5 and 1.7–2.0, respectively; Bourdon et al., 1998) but larger than those in Hawaii (1.0 –1.3 and 1.3–1.8, respectively; Sims et al., 1999; Pietruszka et al., 2001). These differences are wholly consistent with the greater solid upwelling velocities beneath Hawaii relative to the Canaries and the Comores. Basaltic samples from Iceland range from (230Th)/(238U) of 1.1 to 1.3 (Kokfelt et al., in press) and (231Pa)/(235U) of 1.2 to 2.0 (Pickett and Murrell, 1997). Although solid upwelling velocities beneath Iceland are likely to be larger than the Canaries (Sleep, 1990) predicting lower 231Pa excess, maximum 231Pa excesses are similar in Iceland, the Canaries, and the Comores. This is explained by the thinner lithosphere in Iceland producing a longer melting column allowing (231Pa)/(235U) to reach levels similar to the Canaries. Thus 231Pa excess may depend on melting column length as much as solid upwelling rate. A notable feature of (230Th)/(238U) and (231Pa)/(235U) for OIB is the flattening of the correlation at high 230Th excess (Fig. 13). Whereas (230Th)/(238U) and (231Pa)/(235U) form a positive correlation for OIB at low 230Th excess consistent with a progressive decrease in upwelling rate with increasing disequilibria, the flattening of (231Pa)/(235U) can be explained as the lithospheric lid acting to limit the melting column length for 231 Pa ingrowth. In contrast, 230Th excesses do not appear to be limited by the lithosphere; this is consistent with the similarity of (230Th)/(238U) between OIB and MORB (see below). Like (231Pa)/(235U), (226Ra)/(230Th) in the Canaries (up to 1.8) is slightly greater than the Comores (1.1–1.5; Claude-Ivanaj et al., 1998) and Hawaii (1.1–1.4; Sims et al., 1999; Pietruszka et al., 2001). Thus, as expected from its short half-life and susceptibility to lithospheric transport times, there appears to be little relationship between (226Ra)/(230Th) and buoyancy flux beneath different ocean islands. Furthermore, the extreme sensitivity of (226Ra)/(230Th) to relatively small variations in melt porosity may also play a role in its more variable nature. 231

Fig. 12. 1/Th and (230Th)/(238U)o, as a function of (226Ra)/(230Th)o for the basanite-tholeiite sequence of the Timanfaya eruption and a hypothetical lithospheric melt. Consistent with previous interpretations for this eruption (Sigmarsson et al., 1998; Thomas et al., 1999; Reiners, 2002), excellent linear mixing relationships are observed for all samples from basanites to tholeiites. If we assume the lithospheric endmember has (230Th)/(238U)o ⫽ 1.0 (i.e., no garnet present), the (226Ra)/ (230Th)o from the lithosphere must be 2 or greater and the Th concentration must be ⬍ 1 ppm.

basanite. Perhaps one clue to the physical process involved is the large number of mantle xenoliths observed in the initial basanitic stage of the Timanfaya eruption. If a pulse of buoyant plume-derived basanite ascended through an already diffusively fluxed area of lithosphere, significant fracturing of the lithosphere might occur ripping off xenoliths and opening up a more permeable path for the mixed tholeiitic magma. We have shown that the eastern island tholeiites are consistent with creation by mixing between plume basanites and a high silica melt component, produced either by DIA or by incongruent dissolution of orthopyroxene during direct reaction. This explanation differs significantly from previous models for the Timanfaya eruption and, like the others, provides a working hypothesis for explaining this fascinating suite of lavas. Although we have applied this lithospheric melting model to observations from Timanfaya, similar observations of increasing SiO2 and decreasing MgO and trace element concentrations occur in individual eruptions of intraplate volcanoes globally (Reiners, 2002) which could reflect a similar lithospheric melting process. Finally, we note that this interpretation of lithospheric melting differs dramatically from the model of Class and Goldstein (1997) who suggest that posterosional silica-undersaturated lavas with residual amphibole signatures are lithospheric melts. In contrast, we infer that silica-rich basalts from the Canary Islands with signatures of complete fusion of amphibole reflect lithospheric melting.

U-series disequilibria in the Canary Islands

Fig. 13. (231Pa)/(235U) vs. for (230Th)/(238U) for OIB and MORB. OIB systematically have lower 231Pa excess than MORB which is attributed to reduced melting column length due to the presence of a lithospheric lid. The OIB data also show a systematic positive correlation at low excess that is consistent with progressively decreasing upwelling velocity. OIB and MORB have a relatively similar range in 230 Th excess which is attributed to melting garnet bearing sources at depths ⬎ 100 km in both tectonic settings. 230Th excess generation terminates shallower than 100 km in both settings because of: (1) a lithosphere lid (OIB); or (2) exhaustion of garnet (MORB). The Volcan de Corona samples (open symbols) have no excess 231Pa and define an anomalous behavior not seen in any other samples globally. Data from Sims et al. (1995, 1999), Bourdon et al. (1998), Pickett and Murrell (1997), Lundstrom et al. (1995, 1998, 1999), and Goldstein et al. (1991, 1993).

Comparison of U-series disequilibria provides insight into the melting processes between OIB and MORB. Although magmas from these settings are generally considered to result from very different degrees and depths of melting, the range in (230Th)/(238U) is surprisingly similar between the two. Midocean ridge basalts young enough to retain excess 226Ra have (230Th)/(238U) ranging from 0.9 to 1.35 while (230Th)/(238U) in OIB ranges between 1.0 and 1.6 with most samples having ⬍ 1.4. In contrast, (231Pa)/(235U) in MORB range from 2 to 3.5, a factor of 2 to 3 larger than in OIB (Fig. 13). Similarly, 226Ra excesses in MORB are many times larger than those in OIB with (226Ra)/(230Th) in zero age MORB reaching values of 2.5 up to 4, compared with 1.1 to 1.8 in OIB. Lundstrom et al. (2000) argue that the large excesses of (231Pa)/(235U) and (226Ra)/(230Th) in MORB reflect reactive porous flow ascent of melts through the majority of the ridge melting column. The observed difference between OIB and MORB in these systems is explained by the role of the lithosphere. Asthenospheric melting to shallow depths allows MORB to undergo continuous reaction and re-equilibration maximizing residence time differences. Furthermore, the shallow final depth of melting of MORB results in short transport times once melt-solid equilibrium is lost. In contrast, OIB cannot undergo a vigorous reactive flow process at shallow depths because of the lithosphere and the transport time from the final depth of melt-solid equilibrium to the surface may be significant. Although some Canary Island basalts contain a lithosphere contribution (the “lithospheric trend” of Fig. 5), this process probably has a relatively small effect on 231Pa and 226 Ra excesses in most OIB. Thus, the differences in these two systems between OIB and MORB are fully consistent with the

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role of lithosphere in minimizing OIB melting at shallow depths. On the other hand, the similarity of (230Th)/(238U) between OIB and MORB is consistent with previous suggestions about the source of 230Th excess in MORB (Ben Othman and Allegre, 1990; Lundstrom et al., 2000). If 230Th excesses in MORB simply reflect melting of small-scale garnet-bearing heterogeneities which produce melts that mix with and overwhelm the signature of melting at shallow depths (Lundstrom et al., 2000), then it is reasonable that OIB and the enriched MORB source may have similar initial depths of melting. The long half-life of 230 Th results in (230Th)/(238U) being sensitive to melting at greater depths than those recorded by 231Pa. The similarity of (230Th)/(238U) in MORB and OIB may reflect the fact that the melting process creating 230Th excess in both settings occurs at depths much greater than the typical mature lithosphere-asthenosphere boundary (⬎100 km). Such a conclusion might imply that 230Th excesses are generated at the initiation of melting by small amounts of volatiles as has been inferred from the seismic velocity structure beneath the East Pacific Rise (Toomey et al., 1998). 6. CONCLUSIONS

Based on the combination of major element compositional changes, trace element geochemistry and U-series and radiogenic isotope data, we suggest a two-stage process for the evolution of magmatism in the Canary Islands that relates to the slow absolute plate motion of the African lithosphere (Fig. 14). In all islands, plume-derived magmas are able to ascend through the lithosphere and erupt primitive basanites. The two-stage process can be seen by comparison of volcanic rocks erupted on the younger western islands (stage 1) and older eastern islands (stage 2). During stage 1, plume-derived basanites ascend through cold, thick lithosphere; some of these differentiate toward phonolitic compositions by crystallizing amphibole-rich veins and cumulates in the lithosphere, thus changing the concentrations and ratios of certain incompatible elements. During this stage, plume-derived melts also form a variety of intrusive bodies within the lithosphere. In the second stage (now observable on the eastern islands), repeated passage of basanites through time leads to substantial heating and erosion of the lithosphere, fundamentally altering the nature of melting. The passage of basanites causes melting of the shallow lithosphere through either direct basanite-peridotite reaction or by the diffusive infiltration of alkali process. Lithospheric melts are contaminated by small amounts of melt from residual deposits of stage 1, either amphibole-rich cumulates and migmatites having extreme ratios of (K, Nb, Ta, Ba, Rb)/(Th, U, LREE). The lithospheric melts mix with plume-derived basanite to form primitive tholeiites observed on the eastern islands. Acknowledgments—Hans-Ulrich Schmincke, Dave Graham, Thor Hansteen, Tim Elliot, James White and particularly Liz Widom are gratefully acknowledged for their assistance in collecting and/or providing samples used in this study. We thank Dieter Garbe-Scho¨ nberg and Thomas Arpp (University of Kiel) for carrying out the ICP-MS analyses and Kerstin Wolff (GEOMAR) for the XRF analyses. Reviews by Simon Turner, Chris Hawkesworth, Olgeir Sigmarsson, Ken Sims, and especially Fred Frey led to substantial improvements. Financial support for this project was provided by the National Science Foundation

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Fig. 14. Cartoon depicting the evolution of plume and lithospheric magmas in the Canary Islands. In stage 1, witnessed presently in the western islands, plume magmas ascend through and interact with the cold lithosphere. This results in crystallizing amphibole veins or creating migmatites (leucocratic veins) in basal complex wall rocks (both result in a high Nb/U component deposited in the lithosphere). Sustained and repeated passage of plume magmas leads to lithospheric erosion beneath the eastern islands providing for two other processes affecting magmatic diversity during stage 2. First, a longer plume melting column results from upwelling to shallower depths, producing alkali basalts with high 230Th excess such as those from Volcan de Corona. Second, ascending plume basanites interact with the lithospheric mantle in two ways: (1) they assimilate amphibole veins or migmatites created in stage 1 of island evolution imparting a high Nb/U signature and (2) they react with the mantle lithosphere directly or diffusively in a process which adds SiO2 and dilutes incompatible elements. This latter process results in silica enrichment creating the tholeiites observed in the Timanfaya eruption.

(EAR-9105113; JBG and KH) and the Deutsche Forschungsgemeinschaft (DFG Project HO1833/1). Associate editor: F. Frey

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U-series disequilibria in the Canary Islands 1585; 1646 –sample LP43e from Tim Elliot; 1677; 1712; 1949 –samples LP124793, LP124794 from Hans Schmincke; and 1971–sample LP71–7L from Liz Widom) occurring along the Cumbre Vieja rift system extending south of the Taburiente shield volcano. Volcanism along this rift ranges from basanite to phonolite with mafic rocks being the most dominant. Historic eruptives range from basanite to phonotephrite with single eruptions (e.g., 1949) covering the entire compositional range and erupting from multiple vents often simultaneously. Fluid inclusion studies show that the evolved volcanic rocks from the 1949 eruption formed at both crustal and mantle depths (Klu¨ gel et al., 2000). The flow forming Punta Blanco (ELP7) was sampled on the ring road and must be younger than 18,000 ⫾ 3000 yr, since this is the age of the underlying flow (Carracedo et al., 1996). Olivine basalt sample LP261925A from Hans Schmincke comes from the top of the Cumbre Vieja near Hoyo Negro and was mapped as subhistoric (Carracedo, 1996). Phonotephrite sample WHJLP.92.11 from James White is a volcanic bomb from the Nambroque center dated at 2310 ⫾ 150 yr by 14 C (Carracedo, unpublished data, 2002). Xenoliths are common in the historic eruptives consisting of (1) cumulates containing varying proportions of clinopyroxene, Fe-Ti oxides, amphibole, phlogopite, apatite and hauyne, (2) MORB type gabbros from the oceanic crust, and (3) harzburgitic to lherzolitic mantle xenoliths which in some cases contain amphibole, phlogopite, and apatite. La Gomera is the third westernmost island. It consists of a Miocene (⬃7–12 Ma) shield cycle and Pliocene (⬃3–5 Ma) posterosional cycle (van den Bogaard, unpublished data). Although the shield stage lavas are primarily basanitic, the basal complex (intrusive interior of the volcano exposed through erosion) contains gabbros through phonolitic to syenitic intrusives. The evolved intrusives belong to the later stages of evolution of the shield and posterosional cycles. The Pliocene posterosional cycle ranges from basanite to phonolite-trachyte with evolved volcanic rocks being common in the later part of this volcanic stage (Hoernle, unpublished data), similar to what is observed on Gran Canaria (see below). Tenerife, the central island in the Canary chain, is the largest of the Canary Islands. Three basaltic massifs (Anaga, NE; Teno, NW; and Roque del Conde, S) occur on the three corners of Tenerife and are interpreted to represent three distinct shield volcanoes with ages ranging from ⬃3.5 to 8.5 Ma (Ancochea et al., 1990). The center of the island consists of the younger Las Can˜ ados composite volcano and the Dorsal Ridge connecting this volcano to the Anaga Massif. The Pico Viejo and Teide volcanoes are located within and postdate the Las Can˜ ados Caldera formed at ⬃0.6 Ma. The central volcanoes and Dorsal Ridge belong to the late stage of volcanism of the shield cycle of volcanism on Tenerife. The composition of volcanic rocks from the central volcanoes ranges from basanite-alkali basalt to trachyte-phonolite, including large volumes of highly evolved trachy-phonolite. Recent basanitic to trachy-phonolitic eruptions (1430?; 1704; 1705–sample ET6; 1706; 1798; 1909 –sample ET1577) were primarily erupted from fissures on the Los Can˜ ados Volcano and the Dorsal Ridge (Fuster et al., 1968). We also sampled basaltic flows northeast of Puerto de la Cruz (ET2), north of Arafo (ET7), and at Puerto de Guimar (ET8), and a tephriphonolite flow near Montan˜ a de la Cruz de Tea (ET28), all of which were mapped as belonging to the Serie Basaltica Reciente (Fuster et al., 1968). ET8 has (226Ra)/(230Th) ⫽ 1.0 so that it is older than 8000 yr. The ⬃25,000-yr age for these samples is estimated from field observations. Gran Canaria is the third island from the east and can be divided in three cycles of volcanic activity: (1) a Miocene Shield Cycle (⬃8 –15 Ma), (2) Pliocene Posterosional cycle (⬃1.7–5.6 Ma) and (3) a Quaternary Posterosional Cycle (⬍1.3 Ma) (Hoernle and Schmincke, 1993a, 1993b). The Posterosional Cycles have distinct (more depleted) Sr-Pb isotopic compositions from the Shield Cycle on Gran Canaria with posterosional volcanism on all Canary Islands generally having 206 Pb/204Pb ⬍ 19.5 (Sun, 1980; Hoernle et al., 1991; Hoernle and Tilton, 1991; Geldmacher et al., 2001). A given cycle is characterized by an initially mafic stage of volcanism with the degree of silicasaturation increasing with decreasing age. In the case of the Pliocene and Quaternary Posterosional Cycles (the early stage of the Miocene Shield Cycle is not subaerially exposed and has not been sampled), the initial volcanic rocks are highly silica-undersaturated nephelinites and basanites which give way to alkali basalts and a transitional tholeiites.

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Several cubic kilometers of transitional tholeiitic rocks were erupted in a single event at ⬃4.0 Ma which began with the eruption of smaller volumes of basanite and alkali basalt, analogous to the Timanfaya eruption (1730 –1736) on Lanzarote. During the waning stage of the Miocene Shield and Pliocene Posterosional cycles (the Quaternary Posterosional Cycle is still in its initial stage), large volumes of evolved volcanic rocks (hawaiite through trachyte and peralkaline rhyolite and tephrite through phonolite were erupted) with lower volumes of mafic volcanic rocks (alkali basalts, basanites and nephelinites). Although crustal interaction appears to be minor in the mafic volcanic rocks (see discussion in Hoernle, 1998), the moderately alkaline to transitional tholeiitic rocks within both the Miocene Shield and the Pliocene Posterosional Cycle, as well as the evolved volcanic rocks from both cycles, show significant interaction with the lithospheric mantle in their Sr-Nd-Pb isotopic compositions (Hoernle et al., 1991a; Hoernle and Tilton, 1991; Hoernle, 1998; Widom et al., 1999). The extent of the lithospheric interaction, however, was not quantified, due to the uncertainty of the isotopic composition of the enriched lithospheric mantle beneath the eastern Canary Islands. There have been no historical eruptions on Gran Canaria. However, several eruptive centers have been radiocarbon dated using charcoal within or beneath the eruptive products of the vents. Charcoal from Volcan de Arenas (Montan˜ a Negro, sample EGC6) was dated at 3075 ⫾ 50 yr. Charcoal beneath lava from small cinder cone (Montan˜ a Rajada, sample EGC3) northwest of Telde produced an identical age within error of 3080 ⫾ 80 yr (Schmincke, personal communicatino). A similar age was assumed for sample EGC7 from a lava flow from Volcan El Sao, south of Artenara. EGC8 is from a flow from a vent near Caldera de los Marteles believed to be Holocene in age (Schmincke, personal commmunication). Fuerteventura is the second largest island and is located ⬃100 km from the African continental margin. The uplifted basal complex on this island provides a unique glimpse of the deep, internal bowels of a Canary shield volcano. Included in the basal complex are intrusive rocks ranging from ijolite-syenite-carbonatite complexes to nephelinites and basanites through phonolites to alkalic and tholeiitic gabbros. Screens between the intrusives range from late Cretaceous continental rise sedimentary sequences to blocks of ocean crust as well as older volcanic rocks. The oldest dated alkalic intrusive rocks in the basal complex may be as old as Late Cretaceous (Le Bas et al., 1986) with syenites (consisting of akali feldspar, nepheline, aegirine augite, biotite, sphene, magnetite, apatite and zircon) producing Ar-Ar plateau ages of 63.1 ⫾ 0.8 and 64.2 ⫾ 1.0 Ma (Balogh et al., 1999). It is, however, not clear if these syenites were derived from the Canary hotspot or an earlier magmatic event. In either case this magmatism is likely to have affected the thermal and chemical state of the lithosphere into which it was intruded. The emergence of the basaltic shield above sea level was probably coincident with the emplacement of syenite-ijolite-carbonatite intrusions ⬃22 to 23 Ma ago and the main dike swarm between 20 to 24 Ma (Balogh et al., 1999). The carbonatites are so¨ vites consisting of calcite, sanidine, phlogopite, aegerine apatite, pyrochlore, magnetite and zircon. The nepheline syenites contain sanidine, aegerine augite, nepheline, biotite, magnetite, sphene and pyrochlore. The ijolites, syenites and carbonatites, as well as associated nephelinites, basanites and their differentiates have overlapping Sr-Nd-Pb isotopic compositions, indicating their derivation from a common source (Hoernle and Tilton, 1991; Hoernle et al., 2002). The carbonatites, nephelinites (ijolites are intrusive equivalents) and basanites are interpreted as different degree mantle melts, whereas the syenites represent differentiates of the nephelinites. Leucocratic veins with bulk compositions ranging from almost pure albite to pure orthoclase but with similar Sr-Nd-Pb isotopic compositions (Hoernle and Tilton, 1991) to gabbros, syenites and other basal complex intrusive rocks are interpreted as crystallized lowpressure melts of the intrusive rocks (migmatites). Shield cycle volcanism continued until ⬃13 Ma (Coello et al., 1992). After an ⬃7-Ma volcanic hiatus, a Posterosional Cycle began at ⬃5 Ma and may continue into the Holocene, although no units have yielded Holocene ages or have excess 226Ra. The Posterosional Cycle has more depleted radiogenic isotopic compositions than the Shield Cycle, characterized by less radiogenic Sr and Pb and more radiogenic Nd isotopic composition (Hoernle and Tilton, 1991). The older rocks of the Posterosional Cycle are basanitic to alkali basalt, while the youngest rocks

4176

C. C. Lundstrom, K. Hoernle, and J. Gill

also include transitional tholeiites (e.g., sample EF6 and EF9). Samples EF7, 9 and 12 come from vents that might be Holocene in age from field criteria (Fuster Basaltic Series IV) (Coello et al., 1992; Schmincke, 1976). They range in composition from alkali basalt (EF7) to transitional tholeiite (EF9 and 12). Samples EF1 (alkali basalt) and EF5 (basanite) come from a series of cones in southeastern Fuerteventura that are believed to be Holocene in age (Basaltic Series IV of Fuster et al., 1968). Prehistoric transitional tholeiite sample EF6 is from the coast east of La Oliva and near La Caleta. Prehistoric alkali basalt sample EF10 is from the Los Olivos eruptive center. We assumed a 5000-yr age for all of these samples as a Holocene midpoint. Lanzarote is the easternmost of the Canary Islands. Based on the presently available age data, volcanism on Lanzarote can be divided into a series of cycles or pulses: (1) 15.5 to 12.3 Ma, (2) 10.2 to 8.3 Ma, (3) 6.7 to 5.3 Ma, (4) 3.9 to 3.7 Ma and (5) 1.8 to 0 Ma (Coello et al., 1992). The radiogenic isotope data, in particular Pb, suggest that all of the subaerially exposed volcanism on Lanzarote belongs to the posterosional cycle (Sun, 1980; Hoernle, unpublished data) and therefore that the submarine portion of Lanzarote is significantly older than 16

Ma. The final posterosional cycle on Lanzarote began with basanitic to alkali basaltic melts, e.g., Volcan de Los Helechos and Volcan de Corona (20,000 yr; Carracedo, personal communication, 2002) (samples EL15, 17, 18, 20, 21 and possibly 23). The Timanfaya eruption was one of the most spectacular historic eruptions. It began in September 1730 with the eruption of highly silica-undersaturated basanite (samples EL10, 11), laden with mantle xenoliths. Almost immediately, the composition shifted to alkali basalt (samples EL5, 13), which was followed by tholeiite (samples EL3, 8) after about 6 months. From 1731–1736 primarily tholeiite was erupted, reaching an erupted volume of 2 to 3 km3 (Carracedo et al., 1990, 1992). The products of the Timanfaya eruption covered nearly one third of the island of Lanzarote, which had nearly been eroded to sea level. It is ironic that these historic tholeiites, erupted on what may be the oldest of the Canary Islands, are the most silica-saturated rocks in the entire Canary Islands. The most recent eruptions on Lanzarote occurred in 1824, when basanite was erupted from three separate vents (sample EL1 from the easternmost vent).

U-series disequilibria in the Canary Islands

4177

APPENDIX B Replicate U-series Analyses and Errors

Lanzarote EL3 EL3 EL5 EL5 EL8 EL8 EL17 EL17 EL21 EL21 Fuerteventura EF1 EF1 EF5 EF5 EF6 EF6 EF9 EF9 Gran Canaria EGC3 EGC3 EGC6 EGC6 Tenerife ET2 ET2 ET6 ET6 ET8 ET8 La Palma ELP1 ELP1 LP124793 LP124793 LP124794 LP124794 LP71-7L LP71-7L WHJLP92.11 WHJLP92.11 Hierro EH4 EH4

(230Th)/ (234U)/ Ra (226Ra)/ (238U)/ (230Th)/ (232Th) (232Th) Error (238U) Error (238U) Error (pg/g) (230Th) Error

Th (ppm)

U (ppm)

Th/U

2.159 2.182 3.321 3.389 1.956

0.561

3.849

0.788

0.845 0.850 0.515

3.932 3.989 3.797

0.772 0.761 0.799

4.491

1.185 1.198 1.467

3.789

0.801

3.899

0.778

2.544

0.676

3.763

0.806

4.131 4.074 2.097 2.124 2.513 2.489

0.997

4.141

0.733

0.413 0.413 0.614

5.084 5.142 4.093

0.597 0.590 0.741

7.360 7.408 8.939

1.795 1.830 2.406

4.101 4.049 3.715

4.961

1.270

4.211

0.939 0.944 0.990 0.993 0.948 0.955 1.167 1.167 1.101

0.011 0.013 0.006 0.004 0.003 0.006 0.003 0.008 0.005

1.19

0.017

1.28 1.31 1.19

0.008 0.007 0.009

1.46

0.006

1.42

0.009

1.001 0.003 0.458 0.461 1.005 0.003 0.292 0.289 0.997 0.005 0.995 0.004 1.008 0.007

0.952 0.955 0.923 0.929 0.718 0.712 0.841 0.838

0.005 0.006 0.007 0.011 0.009 0.005 0.005 0.012

1.18

0.007

1.003 0.002

1.26

0.011

1.001 0.015

1.20 1.21 1.14

0.018 0.009 0.010

0.998 0.008 0.997 0.003 1.007 0.018

0.740 0.749 0.817

0.993 0.995 0.999 0.988

0.007 0.001 0.004 0.008

1.34 1.33 1.22

0.012 0.004 0.006

1.000 0.003 1.004 0.002 0.999 0.003 1.078 1.065

3.905

0.777

0.006

1.000 0.002

4.254

0.713

1.31

0.005

1.000 0.005

5.176 5.211

1.372 1.376

3.774 3.787

0.804 0.801

0.002 0.004 0.003 0.008 0.003 0.002

1.12

0.990

0.868 0.873 0.935 0.937 0.862 0.873

1.07 1.09

0.007 0.004

0.997 0.004 1.003 0.006

7.779

2.199 2.199 3.310 3.334 1.589 1.574 1.873 1.857 6.173 6.222

3.537

0.858

1.141 0.004

1.33

0.009

3.479 3.473 3.564 3.580 3.486

0.872 0.874 0.851 0.847 0.870

0.009 0.008 0.008 0.022 0.009

0.953 0.947

0.006 0.006 0.005 0.018 0.005 0.009 0.003 0.006

1.28 1.27 1.30 1.30 1.27

3.185 3.204

1.115 1.109 1.110 1.101 1.104 1.110 1.176 1.177

1.000 0.004 0.997 0.004 1.003 0.002

1.23 1.24

0.005 0.008

0.999 0.003 1.017 1.012 1.004 0.006 1.009 0.009 3.130 1.001 0.003 3.162

0.987 0.988

3.665 3.676

0.828 0.825

1.205 0.005 1.206 0.004

1.46 1.46

0.009 0.009

1.001 0.003 1.000 0.003

5.721 5.689

11.52 11.58 5.661 5.636 6.531 19.66 19.94 3.618 3.632

1.220 1.227 1.389 1.377

0.018 0.011 0.008 0.009

1.074 1.073

0.010 0.005

1.443 1.435

0.016 0.015

1.191 1.202

0.017 0.007

Pa (231Pa)/ (pg/g) (235U) Error

0.520 0.514

1.878 0.031 1.848 0.027