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Keywords:

  • ocean island basalts;
  • uranium-series isotopes

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

Uranium-series isotopes (238U-230Th-226Ra-210Pb), major element, trace element, and Sr-Nd isotopic data are presented for recent (<60 years old) Galápagos archipelago basalts. Volcanic rocks from all centers studied (Fernandina, Cerro Azul, Sierra Negra, and Wolf Volcano) display 230Th excesses (4%–15%) and steep rare earth element (REE) patterns indicative of residual garnet during partial melting of their mantle source. Rare earth element modeling suggests that only a few percent of garnet is involved. Correlations between (238U/232Th), radiogenic isotopes and Nb/Zr ratio suggest that the U/Th ratio of these Galápagos volcanic rocks is primarily controlled by geochemical source variations and not fractionation during partial melting. The lowest (230Th/238U) ratio is not observed at Fernandina (the supposed center of the plume) but at the more geochemically “depleted” Wolf Volcano, further to the north. Small radium excesses are observed for all samples with (226Ra/230Th) ranging from 1.107 to 1.614. The 226Ra-230Th disequilibria do not correlate with other uranium-series parent-daughter nuclide pairs or geochemical data, suggesting modification at shallow levels on timescales relevant to the half-life of 226Ra (1600 years). The combination of 226Ra and 210Pb excesses is inconsistent with interaction of magma with cumulate material unless decoupling of 210Pb (or an intermediate daughter, such as 222Rn) occurs prior to modification of Ra-Th disequilibria. An intriguing correlation of (210Pb/226Ra)0 with Nb/Zr and radiogenic isotopes requires further investigation but suggests possible control via magmatic degassing and accumulation that may somehow be related to source heterogeneities.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

Uranium-series (U-series) disequilibria in ocean island basalts (OIB) can provide unique insights into melt generation, transport, evolution and degassing over process-relevant timescales [e.g., Condomines et al., 1988; Turner et al., 1997; Bourdon and Sims, 2003; Saal and Van Orman, 2004; Peate and Hawkesworth, 2005; Bourdon et al., 2005; Sims et al., 2008a; Prytulak and Elliot, 2009]. Prior to melting, the solid mantle can be assumed to be an undisturbed system (i.e., it has not experienced chemical fractionation over the last 350 ka) and consequently, all parent and daughter nuclide pairs are in a steady state of radioactive (“secular”) equilibrium. Following fractionation of parent-daughter nuclides (e.g., via partial melting), return to secular equilibrium is dictated by the half-lives of the daughter nuclides. The wide range of parent-daughter pairs within the 238U decay chain affords constraints across a variety of timescales: 238U-230Th (380 Ka), 230Th-226Ra (8 Ka) and 226Ra-210Pb (100a), though very few studies of OIB have investigated all of these systems. Previous U-series studies suggest that melting beneath ocean islands is controlled primarily by variations in solid mantle upwelling rate [e.g., Stracke et al., 2006] evidenced by radial increases (from the inferred plume center) in (230Th/238U) which are observed at several locations e.g., Hawaii [Sims et al., 1999] and Iceland [Kokfelt et al., 2003]. However, compositional source heterogeneities (lithology, mineralogy) and differential extents of melting, can complicate the interpretation of U-series data in oceanic island settings [e.g., Bourdon et al., 2005].

A selection of young (<60 years old) basaltic lava and tephra samples from Fernandina, Cerro Azul, Sierra Negra and Wolf (on Isabela) Volcanoes in the Galápagos Islands were selected for major element, trace element, and Sr-Nd and U-series (238U-230Th-226Ra-210Pb) isotopic analysis (see Appendix A for details of the analytical techniques). These data form the first published U-series study of the Galápagos archipelago and provide insights into the current dynamic state of the plume. Other regional U-series data has been published for the Galápagos Spreading Centre (GSC) and off rift seamounts [Kokfelt et al., 2005]. The young age of the samples also permits determination of the shorter-lived 210Pb nuclide (22 year half-life) for which there are few published data from OIB.

2. Geological Setting

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

The Galápagos archipelago, located on the Nazca Plate 1000 km west of South America (Figure 1), is believed to be the surface expression of a mantle plume [e.g., Feighner and Richards, 1994; Kurz and Geist, 1999; Toomey et al., 2002; Hooft et al., 2003; Villagómez et al., 2007]. 3He/4He data [Kurz and Geist, 1999] and determinations on the thickness of the mantle transition zone beneath the Galápagos [Hooft et al., 2003] suggest that the center of the plume is located directly beneath Fernandina, or ∼40 km to the southwest of the island. However, surface and body wave tomography suggests that the plume may not take a direct path to the base of the lithosphere, but that it is inclined to the northeast at depths less than 150 km [Toomey et al., 2002; Villagómez et al., 2007]. Unlike Hawaii, there is no simple age progression to the islands, although the age of volcanism generally increases away from Fernandina in the direction of plate motion [Sinton et al., 1996].

image

Figure 1. Tectonic setting of the Galápagos archipelago. Arrows indicate the relative movement of the Cocos and Nazca Plates. Volcanoes sampled for this study are shown in bold: W, Wolf; D, Darwin; A, Alcedo; SN, Sierra Negra; CA, Cerro Azul; F, Fernandina. GSC, Galápagos Spreading Centre.

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Geographical geochemical variations in Galápagos archipelago volcanic rocks are well documented and expose a geochemically and isotopically “enriched” east facing horseshoe-shaped region with the most depleted geochemical signatures located in the center of the archipelago [e.g., Blichert-Toft and White, 2001; Geist et al., 1988; Harpp and White, 2001; Hoernle et al., 2000; White et al., 1993]. Four isotopically distinct mantle components are proposed to contribute to magmatism of the islands: three heterogeneous enriched components of the plume, and a depleted mantle component. The latter has been interpreted as either entrained upper mantle [Blichert-Toft and White, 2001; Geist et al., 1988; Harpp and White, 2001; White et al., 1993] or is attributed to be an inherent part of the plume itself [Hoernle et al., 2000]. Combined trace element geochemistry and U-series data from the GSC indicate solid mantle inflow of Galápagos plume-derived, enriched material into the adjacent ridge system [Kokfelt et al., 2005].

3. Results

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

3.1. Major Element, Trace Element, and Radiogenic Isotope Data

The new major element, trace element and radiogenic data presented here (Table 1) are fully consistent with previously published data [e.g., Geist et al., 1988; White et al., 1993; Harpp and White, 2001; Naumann et al., 2002; Geist et al., 2005]. Chondrite-normalized rare earth element (REE) patterns of the basalts (Figure 2) show a significant enrichment of light REE relative to heavy REE with a relatively constant, steep slope, typical of samples from the western Galápagos region [e.g., White et al., 1993; Harpp and White, 2001]. Wolf Volcano is less enriched than the other three volcanoes, possessing a flatter light REE trend (Figure 2) and the most primitive Sr-Nd isotope ratios (0.702837 and 0.513022, respectively; Table 1). A small negative Eu anomaly is observed in both of the Sierra Negra samples.

image

Figure 2. Chondrite-normalized REE diagram for Galápagos archipelago volcanic rocks. Chondrite normalizing values taken from McDonough and Sun [1995].

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Table 1. Major Element, Trace Element, and Sr-Nd Isotope Data of Galápagos Volcanic Rocks
 W953SN0510SN9121CA9804FE0101FE0103FE0104FE0105FE0602
IslandIsabelaIsabelaIsabelaIsabelaFernandinaFernandinaFernandinaFernandinaFernandina
VolcanoWolfSierra NegraSierra NegraCerro AzulFernandinaFernandinaFernandinaFernandinaFernandina
Typetephralavatephralavalavascorialavalavalava
Eruption age198220051979199819951988195819822005
Collection age1995at eruption1991at eruption20012001200120012006
 
SiO248.2447.4847.2347.7048.2348.0248.1047.7547.82
Al2O315.4312.9712.8315.3416.9514.7614.0014.8813.81
Fe2O311.7015.4315.6011.5111.6812.9113.5612.5013.25
MgO6.375.285.147.645.875.956.125.796.41
CaO11.699.799.6512.1511.3911.0110.6911.0410.98
Na2O3.023.102.962.692.732.932.892.912.80
K2O0.420.600.590.510.450.520.510.500.47
TiO22.764.094.132.333.003.483.563.343.34
MnO0.160.210.210.160.160.180.190.170.18
P2O50.330.420.420.270.350.390.400.380.36
LOI−0.64−0.94−0.24−0.73−0.55−0.58−0.54−0.65−1.01
Total99.4798.4298.5299.57100.2599.5699.4698.6298.41
 
Sc37.636.037.135.831.735.239.432.440.2
V287430437328318376385369385
Cr1277985276205134178155161
Ni70.047.446.514964.058.350.858.159.4
Cu121151140119101130120121126
Zn87.4149151103110127133215128
Ga18.422.923.119.322.823.123.022.822.4
Rb4.8611.311.99.376.778.459.056.648.41
Sr376299300353374348336338325
Y31.642.443.026.629.534.537.031.835.2
Zr194248250156192224219219209
Nb13.834.634.823.123.226.926.326.124.8
Cs0.050.120.140.110.080.100.100.070.10
Ba60.6138140134104115117111107
 
La13.121.922.016.315.217.617.416.816.4
Ce33.951.350.936.736.842.741.940.939.5
Pr4.967.067.024.905.165.965.915.715.56
Nd22.530.830.620.922.826.326.125.124.6
Sm5.897.677.615.015.686.556.616.266.25
Eu2.002.502.491.681.922.152.192.072.07
Gd6.378.308.285.266.056.977.166.656.80
Tb0.991.311.300.820.941.091.131.041.07
Dy5.617.467.454.665.356.206.425.826.11
Ho1.101.491.490.931.061.221.271.161.22
Er2.883.993.992.482.803.273.433.073.26
Yb2.343.353.372.072.302.682.842.512.73
Lu0.330.480.480.290.330.380.410.360.39
 
Hf4.205.585.623.484.254.944.834.804.63
Ta0.811.871.871.201.311.551.471.461.40
Pb1.201.532.981.291.191.621.361.401.27
Th0.952.032.081.501.381.591.611.521.48
U0.300.600.600.440.430.510.490.490.46
 
87Sr/86Sr0.7028370.7033760.7033410.7033910.7032440.7032420.7032590.7032480.703265
2SE0.0000050.0000100.0000120.0000080.0000070.0000070.0000090.0000100.000006
143Nd/144Nd0.5130220.5129240.5129280.512923 0.5129320.5129490.5129540.512942
2SE0.0000050.0000100.0000070.000009 0.0000090.0000070.0000130.000009

3.2. Uranium-Series Isotopes

Uranium-series activity ratios and nuclide concentrations of Galápagos volcanic rocks are reported in Tables 2 and 3. The (234U/238U) activity ratios lie within analytical error of secular equilibrium (±8 per mil) indicating minimal posteruption alteration. Furthermore, several samples were collected the day/year they were erupted (Table 1), allowing little time for alteration. U-series disequilibria in the Galápagos volcanic rocks are therefore interpreted in terms of chemical fractionation by magmatic processes (e.g., melt generation and magma transport, magma differentiation and degassing).

Table 2. U-Th-Ra Concentration and Isotope Data of Galápagos Volcanic Rocks
 W953SN0510SN9121CA9804CA085BFE0101FE0103FE0104FE0105FE0602
IslandIsabelaIsabelaIsabelaIsabelaIsabelaFernandinaFernandinaFernandinaFernandinaFernandina
VolcanoWolfSierra NegraSierra NegraCerro AzulCerro AzulFernandinaFernandinaFernandinaFernandinaFernandina
TypeTephralavatephralavalavalavascorialavalavalava
Eruption age1982200519791998200819951988195819822005
Collection age1995at eruption1991at eruption200820012001200120012006
 
U (μg/g)0.3230.5400.5850.4250.2910.4300.4550.4910.4860.450
Th (μg/g)0.9041.7161.9011.4160.9971.3101.3981.5041.4751.398
226Ra (fg/g)133.65245.60259.61219.88163.14170.44248.84196.55195.73190.36
 
(234U/238U)1.0051.0071.0061.0060.9940.9991.0031.0001.0061.007
2SE0.0020.0020.0020.0020.0040.0020.0020.0020.0010.002
(238U/232Th)1.0840.9540.9340.9120.8850.9960.9880.9910.9990.977
2SE0.0020.0020.0020.0020.0110.0030.0020.0030.0030.002
(230Th/232Th)1.1050.9950.9951.0211.0131.0231.0191.0221.0321.018
2SE0.0050.0040.0040.0040.0220.0040.0040.0040.0080.004
(230Th/238U)1.0281.0511.0741.1301.1541.0361.0411.0411.0421.051
2SE0.0050.0050.0050.0050.0280.0050.0050.0050.0080.005
(226Ra/230Th)1.1851.2741.6141.1411.4411.1251.5481.1321.1381.107
2SE0.0120.0080.0490.0140.0110.0100.0700.0100.0210.011
Table 3. 210Pb Data and (210Pb/226Ra) Activity Ratios for Galápagos Volcanic Rocks
SampleaVolcanoTypeEruption AgeCollection AgeYear of Analysis(210Po) dpm/g+/− 1SD(210Pb)0 dpm/g+/− 1SD(210Pb/226Ra)0b+/− 2SD
  • a

    Repeat analysis, rpt.

  • b

    Initial (210Pb/226Ra)0 activity ratios are calculated for the time of eruption.

CA9804Cerro Azullava1998at eruption20080.4830.0120.4830.0171.0010.070
CA085BCerro Azullava2008200820100.4440.0230.4500.0241.2550.136
FE0101Fernandinalava1995200120080.4980.0130.5590.0201.4920.105
FE0103Fernandinascoria1988200120083.0990.0535.2590.1089.6180.588
FE0103 rptFernandinascoria1988200120092.9130.0625.0530.1309.2410.632
FE0103 rptFernandinascoria1988200120103.2310.1505.8180.30710.6401.221
FE0105Fernandinalava1982200120080.5440.0130.6840.0341.5890.161
FE0602Fernandinalava2005200620080.6420.0140.6640.0151.5870.074
SN0510Sierra Negralava2005at eruption20080.5930.0170.5980.0181.1080.068
SN0510 rptSierra Negralava2005at eruption20100.5880.0310.5960.0371.1040.137
SN9121Sierra Negratephra1979199120097.8230.12618.2220.32031.9421.479
SN9121 rptSierra Negratephra1979199120107.8280.34619.3510.90833.9223.344
SN05HallSierra Negralava2005at eruption20090.5690.010    
SN0502Sierra Negratephra2005at eruption20090.5750.009    
SN0511Sierra Negratephra2005at eruption20090.5390.010    
W953Wolftephra1982199520080.6650.0131.1180.0313.8080.217
W953 rptWolftephra1982199520090.5940.0140.9810.0333.3400.228
3.2.1. 238U-230Th

Galápagos volcanic rock (238U/232Th) and (230Th/232Th) activity ratios range from 0.885 to 1.084 and 0.995–1.105, respectively. Common to OIBs, the volcanic rocks from all centers display 230Th excesses ((230Th/238U) > 1) of up to ∼15% (Figure 3). The majority of samples plot close to and within the field for the off-rift seamounts of the GSC [Kokfelt et al., 2005] and define a roughly subparallel array relative to the equiline. However, both samples from Cerro Azul lie within the field for on-rift GSC volcanic rocks analyzed by Kokfelt et al. [2005].

image

Figure 3. (230Th/232Th) versus (238U/232Th) equiline diagram for young Galápagos basalts. Internal 2SE errors are smaller than symbol size for all samples except for one Cerro Azul sample. The equiline, denoting secular equilibrium between (230Th/232Th) and (238U/232Th) activity ratios, is represented by a solid line. Dashed lines are contours of excess 230Th relative to 238U. Galápagos Spreading Centre (GSC) data from Kokfelt et al. [2005].

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(238U/232Th) and (230Th/238U) activity ratios vary systematically with volcano location (plotted as decimal degrees south of Wolf Volcano, Figures 4a and 4d). A good correlation is observed between (238U/232Th) activity ratios and radiogenic isotopes and Nb/Zr ratios (e.g., Figures 4b–4c). A weaker correlation (based on a linear regression) is noted between (230Th/238U) and 87Sr/86Sr and Nb/Zr. (230Th/232Th) does not correlate with indices of differentiation within the eruptive group of Fernandina (for which more than 3 samples were analyzed). However, a general positive correlation with SiO2 is observed when considering the most mafic samples of each center (Figure 4g).

image

Figure 4. (a–i) Selected correlations between (238U/232Th), (230Th/232Th), and (230Th/238U) and volcano locations, Nb/Zr, 87Sr/86Sr, SiO2, and Dy/Yb for Galápagos archipelago volcanic rocks. Errors on isotopic data are internal 2SE. The displayed R2 values show goodness of fit for linear regression of the data. In Figure 4g, R2 was calculated only for the mafic end-members of each volcano as Fernandina (for which more than three samples were analyzed) shows little variation in (230Th/232Th) over a range in SiO2. In Figures 4h and 4i, Cerro Azul has been excluded from the regression due it its anomalous behavior. Goodness of fit is not shown for variations in geochemistry versus volcano location (Figures 4a and 4d), as the volcanoes do not lie perfectly along a north-south line relative to one another.

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3.2.2. 230Th-226Ra

All the Galápagos samples have excess Ra with (226Ra/230Th) ranging from 1.107 to 1.614. No simple relationship is observed between (226Ra/230Th) and (230Th/238U) (Figure 5a), or between (226Ra/230Th) and SiO2, radiogenic isotopes, Ba/Th, Nb/Zr, Dy/Yb or geographical location (not shown). The 226Ra excesses suggest that the (230Th/238U) ratios have not undergone significant decay [cf. Saal et al., 2007].

image

Figure 5. (a) (226Ra/230Th) versus (230Th/238U) for Galápagos volcanic rocks. The equiline for (226Ra/230Th) is shown as a solid line. (b) (210Pb/226Ra) versus (226Ra/230Th) for Galápagos volcanic rocks. The two tephra samples with significantly high 210Pb (and 226Ra) excess are marked by black open circles. The inset diagram shows a close-up view of samples with lower 210Pb excess.

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3.2.3. 226Ra-210Pb

210Pb excesses (relative to 226Ra) are observed in the majority of samples (Figure 5b and inset). Significantly high 210Pb excesses ((210Pb/226Ra)0 > 5) are observed for 2 tephra samples from separate volcanic centers (Figure 5b and Table 3). When the 2 tephras with high 210Pb excess (and high 226Ra excess) are excluded, (210Pb/226Ra)0 exhibits the same relationship with geographic location as (238U/232Th), (230Th/238U), 87Sr/86Sr, and Nb/Zr; 210Pb excess decreases from Wolf to Cerro Azul and is relatively constant within any one volcano (Figures 6a and 6b). No correlation is observed between (210Pb/226Ra)0 and eruption year, indices of differentiation or 3He/4He isotope ratios (not shown) in the limited data set. The reproducibility of 210Pb data is presented in Table 3.

image

Figure 6. (a) (210Pb/226Ra)0 versus volcano latitude (given in decimal degrees south of Wolf Volcano) for Galápagos volcanic rocks (excluding the two tephra samples with high Pb concentrations and (210Pb/226Ra)0 > 4). (b) (210Pb/226Ra)0 versus 87Sr/86Sr for Galápagos volcanic rocks. Inset diagrams shows (210Pb/226Ra)0 versus Nb/Zr. The displayed R2 values show goodness of fit for linear regression of the data.

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4. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

The data presented here sample only a fraction of the total magmatism in the Galápagos islands, but do offer a current and dynamic picture of the plume and cover much of the compositional spectrum observed in the islands, from depleted (Wolf) to enriched (Sierra Negra/Cerro Azul). The restricted nature of the data set necessarily places limitations on the depth of analysis and absolute model solutions. However, it is noted that Wolf Volcano, Fernandina, and Sierra Negra are each described as geochemically “monotonous” (internally) [e.g., Geist and Teasdale, 2001; Geist et al., 2005], supporting the attempts at spatial investigations using a restricted data set.

4.1. Controls on U-Th Disequilibria

The 230Th excesses observed in most OIB (and MORB) are usually attributed to the presence of residual garnet during mantle melting [e.g., Beattie, 1993; Prytulak and Elliot, 2009]. However, it is noted that high-pressure residual clinopyroxene may also create significant 230Th excess in the melt [Wood et al., 1999]. The 230Th excesses and steep REE patterns of Galápagos volcanic rocks (Figure 2) provide support for a garnet-bearing source residue beneath the archipelago. REE constraints on the degree of melting (La/Yb) and amount of residual garnet (Tb/Yb) are shown in Figure 7. Assuming a similar starting composition for all volcanoes, partial melting of a primitive mantle (garnet-peridotite) source requires only a small amount of residual garnet (2%–4%) and small degree of melting (∼3%–5%) to fit the data (noting that the magma source for Wolf Volcano may be more depleted than primitive mantle). These values lie within the calculated range for the degree of partial melting (2.5% to ∼10%) and amount of residual garnet (2%–9%) proposed by Harpp and White [2001] for the same volcanic centers, and the small degree of melting of MORB-source with ∼5% garnet proposed by Geist et al. [2005] for Wolf Volcano. The tight cluster of the Galápagos data on La/Yb-Tb/Yb suggests that despite source compositional heterogeneities (see below), relatively similar degrees of melting and modes of residual garnet are involved at each volcano.

image

Figure 7. La/Yb-Ty/Yb plot showing partial melting (2%–10%) of a primitive mantle source (composition from Sun and McDonough [1989]) using the partition coefficients and modal proportions for garnet-peridotite given by McKenzie and O'Nions [1991]. Varying proportions of residual garnet (1%–5%) are shown. Data symbols as in Figure 2.

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Correlations of (238U/232Th) and (230Th/238U) activity ratios with Nb/Zr (Figures 4c and 4f) and long-lived radiogenic isotopes (Figures 4b and 4e) may be accidental (if the fundamental behavior of each volcano is unique), but suggest that mantle source heterogeneity affects melting processes and the generation of U-series disequilibria in the Galápagos volcanic rocks. A relationship between Sr-Nd-Pb isotopes, element ratios and U-series disequilibria is also reported for GSC samples and is interpreted as mixing between an enriched (Central Galápagos plume) and depleted (upper mantle and/or Eastern Galápagos plume) component [Kokfelt et al., 2005]. The significantly more primitive Sr-Nd isotope ratios, lower LREE/MREE ratio, lower Nb/Zr (but still > MORB), and lower Th/U ratios of Wolf Volcano in the north of Isabela island may signify a larger contribution of depleted upper mantle/plume component compared to volcanoes further south [e.g., Harpp and White, 2001; Geist et al., 2005]. The question then arises whether the U-series disequilibria yield information on melting dynamics of the plume (e.g., melting rate and upwelling velocity), are primarily controlled by source heterogeneity and mixing, or a combination of the two.

Based on geophysical [Hooft et al., 2003] and He isotope [Kurz and Geist, 1999] data, the center of the Galápagos plume is thought to lie beneath or slightly southwest of Fernandina. Therefore, assuming similar source fertility, if melting rate were the primary control on (230Th/238U) disequilibria in Galápagos volcanic rocks, we might expect the highest melting rate, fastest upwelling and therefore, least disequilibria (lowest 230Th excess) at Fernandina. However, this is not what is observed; the lowest excess is observed at Wolf Volcano, north of Fernandina and toward the margin of the geochemically enriched plume trace. On the other hand, at 120–150 km depth the focus of a low velocity anomaly (interpreted as the locus of the upwelling plume) is centered beneath northern Isabela, i.e., Wolf Volcano [Villagómez et al., 2007]. Between 150 and 100 km depth, the plume conduit inclines northward and is located offshore between Wolf Volcano (Isabela) and Pinta [Toomey et al., 2002; Villagómez et al., 2007]. This depth range is compatible with melting of fertile/depleted garnet peridotite [Robinson and Wood, 1998], therefore it is actually reasonable to expect the fastest upwelling rate, and therefore, the lowest 230Th excess at Wolf Volcano.

Variations in the length of the melting column may contribute to the spatial variations observed in U-Th disequilibria. A longer melting column (corresponding to either a higher mantle temperature and/or more fertile source) should correlate with increased residence time in the melting column, and therefore higher 230Th excess. S-wave velocity data suggests that the calculated variation in lithospheric thickness beneath the four volcanoes of this study is relatively small (50–56 km, [Villagómez et al., 2007]) and REE inversion modeling and parameterization suggest that the top of the melting column is between 57 and 58 km beneath all the volcanoes of interest [Gibson and Geist, 2010]. Therefore, the effects of lithospheric thickness variation [Feighner and Richards, 1994; Villagómez et al., 2007] do not seem to be a significant control on the melting regime in this part of the archipelago. Consequently, in order to create a longer melting column, variable depths to the base of the melting column would be required. Assuming a similar source lithology and mineralogy, the volcano with the highest MREE/HREE ratio (indicating the greatest garnet involvement) may correspond to the deepest melt and therefore, the longest melting column. As a result, it would have the longest time for in-growth and so the largest 230Th excess [e.g., Bourdon and Sims, 2003]. However, this is not the case in the Galápagos rocks. We see an inverse relationship between Tb/Yb and (230Th/238U) activity ratios (cf. GSC lavas where a positive correlation is observed [Kokfelt et al., 2005]). Note that the trace element variations induced by differences in degree/depth of melting are likely to be small compared to the large variability in Nb/Zr ratios and radiogenic isotope ratios inherited from the source.

The extent to which variations in source lithology, such as varying proportions of mafic pyroxenite/eclogite and garnet peridotite contribute to the extent of U-series disequilibria is not easy to assess (e.g., see the discussion by Stracke et al. [2006] and references therein). In a simple scenario, assuming comparable trace element fractionations (bulk DU/DTh ratios) for pyroxenite/eclogite and garnet peridotite during partial melting [Pertermann et al., 2004], a greater contribution of eclogite (and possibly pyroxenite) should correspond with relatively lower (230Th/238U) activity ratios (due to their higher average melt productivity) and generally higher SiO2 contents [Yaxley and Green, 1998; Walter, 1998; Pertermann and Hirschmann, 2003a; Hirschmann et al., 2003]. This general trend is observed in our limited Galápagos data set (Figure 4h), although we acknowledge that small variations in SiO2 can be caused by variations in pressure of melting. However, if eclogite/pyroxenite represent previously recycled oceanic crust or metasomatized material, their 87Sr/86Sr isotopic ratios would be expected to be higher than that of “normal” mantle peridotite. Clearly, this does not fit with the observed trend of increasing (230Th/238U) with increasing 87Sr/86Sr isotope ratio (Figure 4e). Also, the REE models (Figure 7) require only a small amount of residual garnet (2%–4%) to fit the Galápagos data, and may preclude an important role for eclogite/garnet-pyroxenite in the source, as these would be expected to have a significantly higher garnet mode. Furthermore, the highest SiO2 content in the new data set is still relatively low (48 wt%) compared to that expected in basaltic eclogite melts (>57 wt% SiO2 at ∼3 GPa [Yaxley and Green, 1998; Pertermann and Hirschmann, 2003a]).

The preceding discussion, argues against a significant contribution from eclogite in the source. Nevertheless, melting of garnet-pyroxenite has been inferred to be able to create 230Th excesses [Pertermann and Hirschmann, 2003b] and therefore, as an alternative interpretation, the Sierra Negra samples, which display generally the lowest SiO2, highest Sr isotope ratios, lowest U/Th ratio and highest 230Th excess, could be interpreted to reflect a higher contribution of a low-SiO2 pyroxenite melt, produced deeper (∼35–50 km) than that of peridotite (due to its lower solidus temperature [Pertermann and Hirschmann, 2003b]), resulting in a greater residence time in the melting column, and therefore higher resultant 230Th excess.

In summary, the correlation of U and Th isotopes with other geochemical parameters indicates a significant link between mantle source heterogeneity and melting dynamics. Based on the above discussion and present knowledge it is difficult to isolate a simple and geochemically consistent explanation for the correlations observed between long-lived radiogenic isotopes, trace element ratios and U-Th disequilibria. This is, perhaps, not surprising considering the range of possible source component and process combinations and the limited nature of the data set.

4.2. Observations From Shorter-Lived U-Series Isotopes

In contrast to U-Th disequilibria, the 226Ra excesses in Galápagos volcanic rocks do not correlate with volcano location, long-lived radiogenic isotopes, trace element ratios or other U-series activity ratios (e.g., Figure 5). If U-Th disequilibria is primarily controlled by source heterogeneity, this may suggest that storage and shallow-level effects (e.g., interaction with cumulates [Saal and Van Orman, 2004]) overprint those of the source on a timescale relative to the half-life of 226Ra (1599 years). Intriguingly, and to our knowledge for the first time in OIBs, initial (210Pb/226Ra) activity ratios of Galápagos volcanic rocks appear to correlate with volcano location, Nb/Zr and radiogenic isotopes (Figure 6). The apparent correlation of short-lived (210Pb/226Ra)0 disequilibria with long-lived Sr-Nd isotope variation is difficult to explain as it is usually assumed to reflect fractionation of 210Pb (or 222Rn) from the 226Ra parent within the last 100 years.

(210Pb/226Ra)0 activity ratios may be affected by several processes, such as: partial melting, sulphide fractionation, plagioclase accumulation, magma degassing or interaction with crustal cumulates [e.g., Rubin et al., 2005; Van Orman and Saal, 2009; Berlo and Turner, 2010, and references therein; Condomines et al., 2010]. 210Pb excesses are relatively uncommon in OIBs [Berlo and Turner, 2010]. 210Pb excesses in other volcanic settings are usually explained by 222Rn (a precursor of 210Pb) gas accumulation and fluxing preceding and/or during eruption [Berlo et al., 2004; Turner et al., 2004; Berlo and Turner, 2010; Condomines et al., 2010; cf. Kayzar et al., 2009]. In the Galápagos data set, the highest 210Pb excesses are found in 2 tephra samples, consistent with these being erupted during more gas rich phases of the eruption, but these samples also have higher (226Ra/230Th), and higher total Pb (Figure 5b and Table 1), suggestive of contamination, for example by Pb sublimates [Gauthier and Condomines, 1999]. Nevertheless, excluding these 2 samples, the correlations between 210Pb excesses and other isotope systems in the Galápagos data suggests that other processes capable of fractionating these systems may play a role.

In a recent review of 210Pb-226Ra disequilibria in volcanic rocks, sulphide fraction and accumulation of 210Pb-rich plagioclase were deemed unlikely protagonists as the primary cause of (210Pb/226Ra)0 variations [Berlo and Turner, 2010]. The latter case is supported by the unrealistic amount of plagioclase crystals required to create the observed excesses (e.g., ∼30% to create (210Pb/226Ra)0 of ∼1.4 [Berlo and Turner, 2010]). Up to 20% plagioclase accumulation is proposed in lavas of Wolf Volcano [Geist et al., 2005] owing to the large quantity of plagioclase phenocrysts in the lavas, their zonation patterns and anomalously high Sr and Eu concentrations in high Al2O3 content (>15 wt%) rocks. Plagioclase accumulation is also thought to be important in controlling major element contents at Fernandina [Allan and Simkin, 2000] where plagioclase phenocrysts occasionally comprise up to 40% of the lava. Despite Al2O3 contents >15 wt% in 3 samples of this study (Table 1), none of the samples show particularly anomalous Eu behavior (Figure 2). For plagioclase accumulation to explain the 210Pb-226Ra disequilibria, the plagioclase phenocrysts would have had to be very young (decades or less), require extreme initial (210Pb/226Ra), and for Wolf Volcano, much more than 20% accumulation of zero age crystals would be required to produce the observed (210Pb/226Ra) of 3.8. Therefore, plagioclase accumulation is not our preferred model to account for (210Pb/226Ra)0 disequilibria in Galápagos volcanic rocks.

Pb is more compatible than Ra in all likely mantle minerals [Blundy and Wood, 2003]. Therefore, it is extremely difficult to explain 210Pb excess by partial melting in the mantle [e.g., Rubin et al., 2005]. This would require a residual phase with DRa > DPb, such as phlogopite, which has so far only been implicated as a possible residual phase at Floreana [Bow and Geist, 1992], southeast of the study region.

Saal and Van Orman [2004] suggest that high 226Ra excesses in oceanic basalts are produced by diffusive exchange between new batches of magma and previously crystallized cumulate minerals (plagioclase and/or clinopyroxene) in the crust However, this process should produce 210Pb deficits, not excesses in the melt and therefore, is inconsistent with the Galápagos Ra-Pb isotope data.

At face value, the correlation of (210Pb/226Ra)0 with Zr/Nb and long-lived radiogenic isotope suggests that 210Pb excess is related to melting and source variations (similar to the conclusion reached by Kurz and Geist [1999] for helium isotopes of Galápagos basalts) rather than degassing and/or shallow level processes. However, as shown above, partial melting of typical mantle mineralogies is not consistent with 210Pb excesses. If 210Pb excesses are related to 222Rn gas accumulation at shallow levels in a closed system model [e.g., Condomines et al., 2010] for example, via the stalling of ascending magma as it encounters a more viscous mush layer, it is not obvious why 210Pb-226Ra disequilibria might correlate with source heterogeneity geochemical indicators. Perhaps volatile content variability is related to source heterogeneity? It is also not clear why the most primitive lavas possess the highest 210Pb excesses. Furthermore, for 210Pb excess to correlate with long-lived geochemical features, it requires decoupling from Ra (the parent nuclide), suggesting possible segregation of Pb or an intermediate daughter (e.g., Rn) from Ra in the melt prior to modification of Ra-Th disequilibria at shallow levels. More data is clearly required to understand the fascinating relationships between long-lived and short-lived U-series isotopes and other geochemical data.

5. Summary and Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

1. 230Th excess, REE patterns and REE/REE ratio modeling are consistent with the production of Galápagos magma by similar, small-degree partial melts of a garnet-bearing mantle source.

2. Correlations between (238U/232Th), 87Sr/86Sr isotope ratio and Nb/Zr suggest that the U/Th ratio of the lavas is controlled by source heterogeneity and not element fractionation during melting. The data are generally inconsistent with a significant role for eclogitic material in the source, but the involvement of garnet-pyroxenite cannot be ruled out.

3. Unlike other plumes such as Hawaii and Iceland, which show radial increases in (230Th/238U) with increasing distance from the plume center, initial U-series data from the Galápagos do not conform to this model if Fernandina is taken to be the locus of mantle upwelling; the lowest (230Th/238U) observed in the present data set is measured at Wolf Volcano (the most geochemically depleted center studied). However, geophysical evidence suggests that at depths consistent with the onset of melting of garnet-bearing source material (e.g., garnet-peridotite) the plume center may indeed be beneath Wolf Volcano and this may provide an explanation for the observed variation in (230Th/238U) ratios.

4. The lack of correlation between (226Ra/230Th) and other geochemical data suggests that (226Ra/230Th) activity ratios are modified at shallow depth, for example via interaction with cumulate material. However, this model is inconsistent with the preservation of 210Pb excess.

5. 210Pb excesses are common in Galápagos volcanic rocks. Despite previously published evidence for plagioclase accumulation in rocks of the same volcanic centers, it is shown that this process is unlikely to be responsible for the 210Pb excesses. More data is required to confirm and investigate the correlation of (210Pb/226Ra)0 with volcano location, radiogenic isotopes and trace element ratios. Considering that 226Ra is parental to 210Pb, these strong correlations suggest decoupling of 210Pb or an intermediate nuclide, such as 222Rn, from 226Ra prior to modification of (226Ra/230Th) activity ratios.

Appendix A:: Analytical Techniques

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

The Galápagos rock samples were crushed between steel hydraulic plates and then ultrasonicated for 20 min in deionized water (3 times) before drying in an oven at 70°C. The samples were reduced to a fine powder using an agate mill at ∼700 RPM for 10–15 min.

Major element contents were determined on glass discs made by the fusion of ∼0.4 g of powdered rock sample and ∼2.7 g of 12:22 flux using standard techniques. Analysis was carried out using the X-ray fluorescence spectrometer (Spectro XLAB 2000) at the University of Wollongong following the methods of Norrish and Chappell [1977]. In-house rock standards were used to calibrate the machine and monitor accuracy and precision during analysis.

For trace element analysis, approximately 0.1 g of sample was digested in an HF-HNO3-HCl mixture. Samples were diluted with 2% HNO3 to a ∼1/1000 solution. A 5 mL aliquot was used for analysis and was spiked with an internal standard containing known amounts of 6Li, 75As, 103Rh, 115In and 209Bi for machine calibration. Samples were analyzed on an Agilent 7500 ICP-MS at the Geochemical Analytical Unit (GAU) at Macquarie University. International reference materials (BIR-1 and BHVO-2) were prepared and analyzed alongside the samples to determine accuracy of the method.

Sr and Nd-isotope cuts were taken from the first Ra cationic column and were prepared and analyzed at the GAU at Macquarie University following the same methods as those described by Heyworth et al. [2007] and Handley et al. [2008]. The Sr and Nd fractions were loaded onto degassed single and double Re filaments, respectively, and analyzed in static mode on a ThermoFinnigan Triton® TIMS. Mass fractionation was corrected for by normalizing Sr to 86Sr/88Sr = 0.1194 and Nd to 146Nd/144Nd = 0.7219. Over the period of study, analyses of NIST SRM-987 gave 87Sr/86Sr of 0.710210 ± 37 (2SD, n = 14) and BHVO-2 gave 143Nd/144Nd of 0.703490 ± 52 (2SD, n = 15) while the JMC Nd standard gave 0.511106 ± 2 (2SD, n = 8) and BHVO-2 yielded 0.512967 ± 6 (2SD, n = 17).

Uranium-series isotopic compositions were determined using the procedure employed at the Macquarie University U-series research laboratory for volcanic rock samples [e.g., Handley et al., 2008; Beier et al., 2010]. Typically, 0.5–1 g of rock powder was spiked with 236U-229Th and 228Ra tracers and digested in mix of HNO3-HCl-HF. Separation of U, Th and Ra followed methods described by Heyworth et al. [2007] and Turner et al. [2007]. Th and U were analyzed on a Nu Instruments® MC-ICP-MS at Macquarie University following the approach described by Dosseto et al. [2006], Heyworth et al. [2007] and Sims et al. [2008b]. Accuracy (<0.3%) and precision (<0.1%) were assessed by regular analyses of the U010 and ThA solution standards. Over the period of study, replicate analyses (n = 5) of TML gave (234U/238U) = 1.001 ± 0.004 (SE), (230Th/232Th) = 1.082 ± 0.003 (SE), and (230Th/238U) = 1.004 ± 0.005 (SE). TML activity ratios are within error of published values [e.g., Sims et al., 2008b] and secular equilibrium for (230Th/238U). Ra analyses were performed on a ThermoFinnigan Triton® TIMS at Macquarie University following the procedures described by Turner et al. [2000]. Samples were loaded onto degassed single Re filaments using a Ta-HF-H3PO4 activator solution [Birck, 1986] and 228Ra/226Ra ratios were measured to a precision typically ∼0.5% in dynamic ion counting mode. Accuracy was assessed via replicate analyses of TML that yielded 226Ra = 3532 fg/g and (226Ra/230Th) = 0.992 ± 0.006 (n = 4). Galápagos (230Th/232Th) and (226Ra/230Th) ratios have not been recalculated for differences in eruption age as all samples were erupted in the last 60 years and therefore posteruption radioactive decay is insignificant compared to the half-life of 230Th (75,690 years) and 226Ra (1599 years).

210Pb procedures are based on those of Reagan et al. [2005] and assume secular equilibrium between 210Pb and 210Po. Two grams of rock powder were spiked with 1g of 209Po tracer and digested in a HF-HCl-HNO3 mixture. Samples were taken up in 1M HCl and loaded onto a 3 ml AG X8 anionic resin column. Most major and trace elements were eluted in 1 N and 0.5 N HCl after which Po was collected by elution of 85 ml of warm 7.5 N HNO3. Po was autoplated onto Ag discs fixed on a magnetic spinner for 6–8 h in 150 ml of warm 0.5 N HCl which contained ∼0.05g ascorbic acid. 210Po and 209Po were counted for approximately a week using an OCTETE Plus® alpha spectrometer at Macquarie University. Several of the samples were analyzed multiple times to check reproducibility (Table 3 of the manuscript). The errors on (210Pb/226Ra)0 are dominated by the analytical error on 210Po and this is significantly magnified by the extrapolation to eruption age in those samples which are older than one half-life (22.6 years). The alpha spectrometer is used exclusively for Po analysis and so background counts were negligible (0.0002 dpm/g). Activity ratios (denoted by brackets) for all U-series data were calculated using the half-lives compiled by Bourdon et al. [2003].

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References

We are grateful to Norman Pearson and Peter Wieland for invaluable assistance during analytical work at Macquarie University. We thank Dennis Geist and William M. White for their reviews and Joel Baker for the editorial handling of the manuscript. The analytical data were obtained using instrumentation funded by DEST Systemic Infrastructure grants, ARC LIEF, NCRIS, industry partners, and Macquarie University. S.T. acknowledges the support of Australian Research Council Federation and Professorial Fellowships. This is contribution 752 from the Australian Research Council National Key Centre for the Geochemical Evolution and Metallogeny of Continents (http://www.gemoc.mq.edu.au).

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Geological Setting
  5. 3. Results
  6. 4. Discussion
  7. 5. Summary and Conclusions
  8. Appendix A:: Analytical Techniques
  9. Acknowledgments
  10. References