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1 doi: /nature ε 183 W values in samples and reference materials Sample analysis yields not only mass-biased corrected 182 W/ 184 W but also 183 W/ 184 W. As an internally normalised stable isotope ratio, we would expect these values to be invariant and the same as the standard, namely ε 183 W = 0. Deviations of ε 183 W from zero thus reflect unexpected aspects of analysis. As shown in Fig. S1, the ε 183 W values in the Isua samples and the reference materials are very close to zero but extend to slightly negative values. Average values of ε 183 W are ± 0.04 (2σ) for Isua and ± 0.02 (2σ) for post-archean silicate samples. In contrast, the average ε 183 W value of the Alfa Aeasar W reference solution is 0.00 ± 0.02 (2σ; Supplementary Information Table S2) suggesting that the offset to negative ε 183 W is restricted to samples processed through the chemical separation process. This view is supported by the finding that (a) the ε 183 W of an aliquot of the Alfa standard solution passed through our chemical separation process is ± 0.02 (2σ m; n=1; Supplementary Information Table S2) and (b) the same magnitude of offset is observed for samples passed through both types of chemical separation (i.e. the acid medium and the alkaline sinter techniques, see Methods section). We can exclude that the negative ε 183 W values are caused by insufficient or erroneous correction of an instrumentally induced mass fractionation, because the ε 182 W and ε 183 W values are negatively correlated (Fig. S1). Clearly the negative ε 183 W cannot be generated by interferences on mass 183 itself, but the trend to negative ε 183 W and positive ε 182 W could possibly be caused by coupled, minor (sub 10ppm) interferences on the even W isotopes. For example, an interference on mass 184 would lower both 183 W/ 184 W and 182 W/ 184 W but if there was an associated interference on mass 182 it could counter-act the effect of the former. Other possibilities can be envisaged. We can exclude that 184 W and 186 W have been affected by isobaric 184 Os and 186 Os isotopes as the intensities for all Os masses monitored are at baseline levels and insufficient to make even these minor perturbations 1

2 Figure S1. ε 182 W versus ε 183 W for quintiple measurements of Isua samples (red symbols) and post-archean samples (black symbols). Error bars as 2σ m. Data from Supplementary Information Table S1. See text for details. Other potential interferences on masses 182 W, 184 W, and 186 W are 91 Zr 2, 92 Zr 2, and 93 Nb 2. We tested the effect of these possible interferences on the ε 182 W and ε 183 W values using a 100ppb Alfa Aesar standard solution doped with Zr and Nb at 100ppb to 10ppm levels. The results demonstrate that there is no correlation between the measured intensity at mass 188 (monitoring 94 Zr 2 ) and ε 183 W (mean ε 182 W = 0.01 ± 0.05; mean ε 183 W = 0.01 ± 0.03; 2σ; n=5). Even for a Zr/W ratio that approached a value of 100, no measurable change in ε 183 W could be observed. For all ε 182 W and ε 183 W values reported here, the Zr/W ratio was less than 0.2 (and Nb/W less than 0.02), as determined on splits of the samples measured on our ThermoFinnigan Element2, thus removing the possibility that Zr or Nb dimers interfer significantly with the even numbered 182 W, 184 W, and 186 W isotopes. This view is supported by the fact that an aliquot of Alfa Aesar solution processed through the anion exchange chemistry also shows a negative ε 183 W value despite containing no Zr. We can also exclude interferences by incompletely separated matrix elements as (a) the Alfa solution processed through the acidic anion exchange chemistry, and (b) samples processed through two entirely different chromatographic exchange chemistries 2

3 (acidic and alkaline) all show a similar level of ε 183 W depletion. If incompletely separated matrix elements caused interferences one would expect the extent of ε 183 W depletion to be different in all three cases. The only common denominator between all three procedures is the type and amount of anion exchange resin (2 and 3ml of AG1x8). If the apparent depletion in ε 183 W is caused by interferences on even mass W isotopes these interferences must have been organic compounds formed by resin breakdown products even after repeated treatment of the samples with a mixture of 15M HNO 3 and 30% H 2 O 2 (m/v). Perhaps more plausibly we noted that recent studies have shown that massindependent isotope fractionation can occur during chemical exchange reactions 1 including liquid chromatography 2. Typical differences in nuclear charge radii between odd and even mass nuclides can lead to mass-independent isotopic fractionations, especially in heavy elements where the effect of nuclear charge radius becomes more important 7. Such a process could explain the deficit in the only stable odd mass W isotope, 183 W, in samples that have been processed through our separation chemistry. In the Methods section, we noted that re-dissolution of W after chemical separation in some cases was incomplete, allowing potentially significant fractionations between sample removed in the liquid phase and the residue. We therefore suggest that the apparent depletion in 183 W was caused by mass-independent fractionation due to nuclear charge effects, which likely occurred during final sample take-up before mass spectrometric analysis. This is also in keeping with the fact that not all samples show anomalous ε 183 W. No nuclear charge radius data for W is currently available that would enable us to directly test this hypothesis. However, if sufficiently precise measurements could be made, this model would predict that samples with anomalous ε 183 W would be the most fractionated relative to the standard. These measurements might be possible using a double spike, but are beyond the scope of this study. A measurement of the amount of W lost during chemical separation might serve as a rough proxy for such fractionation, but unfortunately we do not have sufficient information to be able to check the yields of our different samples and to see how these might correlate with measured ε 183 W. Independently from our study, odd-even isotope separation in W isotopes in MC- ICPMS when using x skimmer cones has recently been described by ref. 3 who 3

4 speculated that this is due to a more efficient neutralisation of 183 W relative to the even W isotopes in the expansion jet behind the sampler cone. Yet importantly, whatever the reason(s) for these small shifts in the 183 W/ 184 W ratio, these effects are small and do not compromise the conclusions of our work as pointed out in the following. 4

5 2. Accuracy of ε 182 W values We correct the ε 182 W for associated perturbed ε 183 W values using the correlations between ε 182 W and ε 183 W values (Fig. S1). Linear fits (York regressions) of ε 182 W and ε 183 W for both datasets (Isua and post-archean samples, respectively) yield regressions with constant slopes. Using these regressions to correct the sample arrays to ε 183 W = 0 yields an average ε 182 W = 0.13 ± 0.04 (2σ) for the Isua dataset and ε 182 W = ± 0.02 for post-archean samples. These values are within error of the uncorrected averages of ε 182 W = 0.18 ± 0.05 and ε 182 W = 0.02 ± 0.05 (2σ) respectively. Regardless of these complications, the Isua data show a well-resolved positive offset of at least 0.06 ε 182 W units compared to post-archean silicate samples. 5

6 3. Description of post-archean samples (oceanic basalts) In addition to the high-precision, quintiple measurements given in Tables 1, S1, and S2, we analysed the W isotopic compositions of a wide range of modern oceanic basalts from the Azores, Cape Verde, La Palma, Hawaii, the Ontong Java Plateau, the Manihiki Plateau, and the Shatsky Rise using one measurement only (i.e. one chemical separation and mass spectrometric analysis). The complete dataset is given in a separate electronic file (Supplementary Information Table S3). The precisions obtained for these single measurements are still better than those of several recent studies of the W isotopic composition of terrestrial samples 4,5. For example, in Fig S2 we compare our data for single measurements of oceanic basalts to the analyses of Isua samples from ref 4. Figure S2. ε 182 W versus ε 183 W for post-archean samples (oceanic basalts) obtained by single measurements (Supplementary Information Table S3; red symbols) compared to recent literature data 4 (black symbols). Error bars are 2σ m. Samples from the Azores comprise basalts from the islands of Sao Miguel (SM) and Pico (P). All six samples are mildly alkalic basalts and cover the entire isotopic range of these islands 6,7. Samples from the Cape Verde include six basalts from the islands of Fogo (FG) and Sao Antao (SA). All samples are young (<2.0 million years old), fresh, in part vesicular basalts and are aphyric to slightly olivine- or pyroxene-phyric. Six samples from La Palma are all <4.0 million years old 8,9 and were taken from the submarine section 6

7 (LP-134), the extinct, subaerial Taburiente volcano (LP-107, LP-113, LP-111) and the historical activity on the Cumbre Vieja (LP-30a, LP-45d). Further geological background as well as supporting geochemcial and isotopic data can be found in ref. 10. The samples from Hawaii comprise the most primitive tholeiitic picrites (Hualalai, H 11; Loihi seamount, LO-02-02, LO-02-04; Mauna Loa, ML-2-50) obtained from M. D. Norman (Australian National University) for which 182 W/ 184 W isotope data and other isotope data has been reported previously Our ε 182 W results for the samples from Hawaii (average -0.02±0.1; 2σ) are in good agreement with the data reported previously for these samples 15 which gave an average ε 182 W of -0.06±0.17 (2σ) for these samples. Our oceanic basalt dataset also includes samples from three Cretaceous Large Igneous Provinces (LIP). Two samples from the Ontong Java Plateau (Pacific Ocean) are tholeiitic basalts recovered during IODP leg 192 and belong to the Kroenke ( A- 006R-06W) and Kwaimbaita ( A-022R-03W) groups The Manihiki Plateau is a ca. 123 million year old submarine plateau, which may have formed as part of the Ontong Java Plateau in the Cretaceous 19,20. We have analysed two samples from the Manihiki plateau recovered during IODP leg 33 ( A W and A- 034R-01W) for which Pb isotopic data has been reported 21. Samples from the Shatsky Rise comprise three samples of the IODP leg All samples ( B W, B-033R-05W, and B-031R-03W) come from the TAMU massif of the Shatsky Rise assumed to represent the site of impingement of the plume head and have a tholeiitic composition. See refs. 22 and 23 for a more detailed description and geochemical data. 7

8 4. Description of the hidden reservoir re-mixing model Conceptual Overview. We wish to test if an appropriate composition of an abyssal, enriched layer can be plausibly generated, such that: a) its formation from a chondritic initial Earth can explain the radiogenic ε 142 Nd observed in 3.8Ga Isua samples 23 and b) its subsequent, partial re-entrainment into the convecting mantle can account for decreases in both ε 142 Nd 23 and ε 182 W (this study) between the 3.8Ga Isua samples and modern silicate Earth. A positive result would add significant support to the hidden reservoir model of ref. 25, whereas a negative result would undermine a powerful line of evidence previously invoked in its defence 26. Admittedly, a failure of our model to reproduce the key observables might simply imply that a more complex model is required. For example, it could be envisaged that a non-chondritic Earth also developed a hidden reservoir. Adding additional degrees of freedom to the model will make obtaining successful solutions more likely but we are solely interested in testing the plausibility of an endmember model, starting with a chondritic composition. Thus we have examined a simple evolution model of a mantle that comprises an accessible, convecting reservoir and a hidden, enriched, abyssal reservoir. Fig S3 provides a cartoon illustrating the reservoirs considered and Fig S4 gives a sketch of example ε 142 Nd and ε 182 W evolution histories. Our model has many similarities to other recent work that quantifies coupled changes in ε 142 Nd and ε 182 W as a result of abyssal reservoir formation 4,5, but a novel aspect of our approach is that we consider the effects of partial remixing of the hidden reservoir to explain a secular decrease in mantle ε 142 Nd and ε 182 W. An initial bulk silicate Earth (BSE) is separated into an enriched and depleted reservoir at a discrete time (T diff ). We use the terminology of ref. 25, referring to the abyssal hidden layer as the Early Enriched Reservoir (EER), and the the accessible Hadean mantle as the Early Depleted Reservoir (EDR), Fig. S3. As in other recent models of consequences of the hidden reservoir for ε 182 W evolution, we assume the timing of this differentiation (T diff ) post-dates core segregation 4,5. If core formation continued after EER formation, arbitrary assumptions would have to be made about the 8

9 fraction of core material that interacted with the different mantle reservoirs. Moreover, the energy associated with accretion during major core growth makes it unlikely that an EER could form and remain isolated before this process had terminated. We assume that the source of the Isua rocks sampled the EDR prior to a fraction of the EER being remixed into EDR to create the current depleted accessible Earth (DAE), see Fig S3. By the time of remixing the short-lived nuclides of interest ( 146 Sm and 182 Hf) have fully decayed (Fig S4). Our DAE represents a mix of depleted mantle (DM) and continental crustal (CC) components. The latter is inferred to have been produced after the mixing of EER with EDR to produce DAE. The BSE at present day is thus divided into a mass fraction X of residual EER and (1-X) of DAE (Fig. S3). The mass of EER (as a fraction of BSE) remixed with the EDR to produce the DAE is dx, such that the BSE was originally divided into (X+dX) EER and (1-X-dX) EDR (Fig. S3). The Sm and Nd distribution between EDR and EER, together with T diff are set by the observed 146,147 Sm- 142,143 Nd systematics of the present day mantle and ε 142 Nd of Isua. By linking the distribution of Hf and W between the reservoirs to that of Sm and Nd via experimentally determined partition coefficients (Table S5) it is then possible to calculate the ε 182 W evolution of the EDR and EER (Fig S4). Strictly, we model the difference in ε 182 W between the EDR and EER, E 182 W EDR-EER, since determining the absolute ε 182 W of EDR requires knowing the timing of T core which we do not constrain (Fig S4). We can calculate a minimum ε 182 W DAE, (i.e. the ε 182 W of the modern day, accessible silicate reservoir), by assuming that ε 182 W of BSE followed a chondritic evolution path. Since BSE has higher Hf/W than chondritic as a result of core extraction this gives a minimum ε 182 W, although if EER formation occurred shortly after core formation the underestimate might not be too extreme (Fig. S4) 9

10 Figure S3. Cartoon showing the reservoirs considered in the model and their changes in size during the Earth history. 10

11 Details of Model The input parameters for the model are given in Table S4. In brief, we use bulk silicate Earth (BSE) values of Nd and Hf concentrations from ref. 30 coupled with chondritic Sm/Nd, 143 Nd/ 144 Nd 27 and 142 Nd/ 144 Nd 24. For less well-constrained parameters we allow a range of values (Table S4) and take a Monte Carlo approach to explore model variability (see below). For example, the loss of W to the core means that the W concentration of BSE is not readily obtained from a chondritic reference and so we use a range of values centred on a recent estimate from ref. 29. We note that by starting our model with a Hf/W ratio appropriate for BSE we implicitly account for this prior core segregation. Using the inputs reported in Table S4 we constrain the ε 182 W evolution of the mantle according to the following equations. The Nd concentration and ε 143 Nd of DAE are calculated from a range of plausible continental crust (CC) and depleted mantle (DM) compositions (Table S4). Namely: Nd DAE = Nd DM (1 F CC ) + Nd CC F CC [1] ε 143 Nd DAE = Nd DM ε143 Nd DM (1 F CC ) + Nd CC ε 143 Nd CC F CC Nd DM (1 F CC ) + Nd CC F CC, [2] where F CC = [3] 27 (1 X) or the weight fraction of DAE present as CC (see Table S4). 11

12 Figure S4. Illustration of example evolution trajectories of ε 142 Nd and ε 182 W for different model reservoirs. The Nd concentration and isotope compositions of EER are then calculated by mass balance. εnd values, as is convention, are referred to chondrite and so ε 143 Nd BSE and ε 142 Nd BSE of the chondritic Earth we model are zero. Unless otherwise indicated epsilon values are reported at present day (T=0). Hence the relevant mass balance equations are: 12

13 Nd EER = Nd Nd BSE DAE (1 X) X [4] ε 143 Nd EER = Nd DAE ε143 Nd DAE (1 X) Nd EER X [5] ε 142 Nd EER = Nd DAE ε142 Nd DAE (1 X) Nd EER X. [6] Similarly, mass balance equations can be written to describe the Nd elemental and isotopic budget of EDR, namely: Nd EDR = Nd BSE Nd EER (X + dx) 1 X dx [7] ε 142 Nd EDR = Nd EER ε142 Nd EER (X + dx). [8] ε 142 Nd EER (1 X dx) Equations [7] and [8] can be solved simultaneously to find Nd EDR and dx, given the ε 142 Nd EDR taken from the average ε 142 Nd from Isua samples (Table S4) reported by ref. 24. The distribution of Hf and W between the reservoirs is then calculated using partition coefficients appropriate for two magma ocean crystallisation scenarios 4. We use the compiled values given in Table S5 to examine both shallow and deep mantle crystallisation. These scenarios respectively correspond to conceptual models of the EER as recycled basaltic crust 25,30 or a residual melt that is denser than the equilibrium solid assemblage in the lower mantle 31. In either case the EER can be thought of as fraction of melt (F melt ) left after crystallisation (or melting) of either upper or lower mantle assemblages. F MELT can be calculated from the already constrained Nd system: 13

14 F MELT = Nd BSE Nd EER DNd Nd EER (1 D Nd ) [9] From equation [9] and the partition coefficients in Table S5 we can calculate: W EER = W BSE D W + (1 D W ) F MELT [10] Hf EER = Hf BSE D Hf + (1 D Hf ) F MELT [11] W EDR = W BSE W EER (X + dx) (1 X dx) [12] Hf EDR = Hf BSE Hf EER (X + dx) (1 X dx) [13] W DAE = W EER dx + W EDR (1 X dx) 1 X [14] Hf DAE = Hf EER dx + Hf EDR (1 X dx). [15] 1 X We can further use the Nd isotope data to constrain the timing of the formation of EER 25,32. The ε 143 Nd data allow us to constrain the Sm/Nd fractionation between EER and the chondritic reference. Using the notation of ref. 33 we thus calculate EER f Sm / Nd ε143 Nd EER Q Nd T diff, [16] 14

15 where T diff is the age of EER formation (as discussed above), Q Nd =10 4 λ 147 Sm 147 Sm 144 Nd 143 Nd 144 Nd CHUR CHUR [17] T = 0 and EER f Sm / Nd Sm Nd = Sm Nd EER CHUR 1. [18] The age of EER formation can be calculated from the 142 Nd system, via the following relationship, based on the notation of ref. 34: ε 142 Nd EER 146 T ' Sm 144 Sm EER q Sm f Sm / Nd e λ146sm (T ' T diff ), [19] where ( 146 Sm/ 144 Sm) T is the initial, chondritic ( 146 Sm/ 144 Sm), i.e. at T or 4567Ga and q Sm = T = 0 Sm 142 Nd CHUR [20] (i.e. the present day value). To solve these two equations we assume an initial T diff of 4537Ma, in keeping with previous estimates from Nd isotope systematics 25, to calculate f EER Sm/Nd and thence T diff. After four iterations the solutions converge. 15

16 Since the Hf and W distribution and age of isolation are now set, we can calculate the W isotope evolution of the reservoirs. Firstly, we calculate the relative difference in W isotopic composition between EER and EDR reservoirs: Ε 182 W EDR EER 182 Hf 180 Hf T ' q EER Hf f EDR EER Hf /W e λ182 Hf (T ' T diff ) [21] where f EDR EER Hf /W = Hf W Hf W EDR EER 1 [22] and 180 Hf EER = W q Hf T = 0 EER [23] It should be noted that ( 180 Hf/ 182 W) is initially unknown, but even using an extreme value for initial chondrite has a negligible influence on the calculation and so we do not iterate this value. Finally the tungsten isotopic composition of the modern, accessible silicate Earth, ε 182 W DAE, can be expressed: ε 182 W DAE = W EDR ε182 W EDR (1 X dx) + W EER (ε 182 W EDR E 182 W EDR EER ) dx W EDR (1 X dx) + W EDR dx [24] To evaluate this expression, however, we require a value of ε 182 W EDR. In our model, this value is left floating as we do not prescribe a time for core formation (Fig S4). However, we can readily obtain a minimum value of ε 182 W DAE by using a chondritic 16

17 starting W isotopic composition for the EDR, namely ε 182 W EDR (T diff ) = ε 182 W CHUR (T diff ). Since Hf/W BSE >Hf/W CHUR as a result of core formation, ε 182 W EDR (T diff ) > ε 182 W CHUR (T diff ) when T diff > T core as is always the case in our model. In the cases where T diff ~ T core the minimum value of ε 182 W DAE is not greatly underestimated. For both scenarios of hidden reservoir formation (i.e. for melt-solid equilibration using partition coefficients for shallow and deep mantle mineralogies) we ran 5 million Monte Carlo simulations of the model using a MatLab programme that allowed each of the variables shown in Table S4 to vary randomly within their permitted range. We excluded results that yielded negative concentrations and T diff <20Ma. The latter was guided by estimates of T core ~30Ma for single stage models 32,34. Since our model assumes T diff > T core setting values of T diff >20Ma represents a generous lower bound. All the calculated minimum ε 182 W DAE for solutions that met these criteria are plotted in Fig. 2 and reported together with associated input parameters and other intermediate, calculated values in Tables S6 (shallow mantle) and S7 (deep mantle). Since all the minimum values thus obtained for ε 182 W DAE are significantly greater than the observed value of zero (Fig. 2), we find no successful solutions to our model. This argues against the plausibility of a hidden reservoir remixing model for explaining the secular decrease in mantle ε 142 Nd and ε 182 W. Qualitatively, the model fails to return plausible ε 182 W because the EER has such fractionated Hf/W. This is a consequence of the large contrast in incompatibility of Hf and W over a wide range of crystallisation scenarios, coupled with the mass balance set by the 147 Sm- 143 Nd system. Once the timing of EER formation is constrained to be early, by the 146 Sm- 142 Nd systematics, a highly radiogenic ε 182 W of the complementary depleted mantle becomes an inevitable consequence. 17

18 Supplementary Information Table S1. ε 182 W and ε 183 W data for samples and NIST SRM 361 steel standard by quintiple measurements. Sample Location Chemistry 1 ε 182 W 2 ±2σ m ε 183 W 2 ±2σ m Orthogneisses _1 Acid _2 Alkaline Average Isua _1 Acid _2 Alkaline Average Isua Metabasalts SM/GR/98/21 Isua Acid SM/GR/98/23_1 Acid SM/GR/98/23_2 Acid SM/GR/98/23_3 Alkaline Average SM/GR/98/23 Isua SM/GR/98/26_1 Acid SM/GR/98/26_2 Acid Average SM/GR/98/26 Isua Amphibolite enclave SM/GR/00/22 Isua Acid Metasediment SM/GR/97/31_1 Alkaline SM/GR/97/31_2 Alkaline Average SM/GR/97/31 Isua Average Isua σ Post-Archean rocks I 21_1 Acid I 21_2 Acid I 21_3 Acid I 21_4 Acid I 21_5 Acid I 21_6 Acid I 21_7 Acid I 21_8 Acid Average I 21 Iceland σ LP68a_1 Acid LP68a_2 Acid LP68a_3 Acid LP68a_4 Acid LP68a_5 Acid LP68a_6 Acid Average LP68a La Palma σ Continued. 18

19 Supplementary Information Table S1. ε 182 W and ε 183 W data for samples and NIST SRM 361 steel standard by quintiple measurements (continued). Sample Location Chemistry 1 ε 182 W 2 ±2σ m ε 183 W 2 ±2σ m LP48c_1 Alkaline LP48c_2 Alkaline LP48c_3 Alkaline Average LP48c La Palma SM-7 Sao Miguel Alkaline BHVO-2_1 Acid BHVO-2_2 Acid Average BHVO-2 Hawaii CS09G01_1 Acid CS09G01_2 Alkaline CS09G01_3 Alkaline Average CS09G01 Guernsey Average post-archean σ NIST SRM 361 SRM 361_1 Acid SRM 361_2 Acid SRM 361_3 Acid SRM 361_4 Acid SRM 361_5 Acid SRM 361_6 Acid Average SRM σ See method section. 2 Average values of five independent analyses. 19

20 Supplementary Information Table S2. ε182w and ε183w data for Alfa Aesar W standard and Alfa Aeasar standard solution pased through chemical separation procedure by quintiple measurements. Run_number ε 182 W ±2σ m ε 183 W ±2σ m Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Alfa_ Average σ Alfa_chem Alfa Aeasr standard solution put through acid chemical separation procedure (see method section) 20

21 Supplementary Information Table S3. ε 182 W and ε 183 W data for oceanic basalts and reference materials by single measurements. Sample Chemistry 1 ε 182 W ±2σ m ε 183 W ±2σ m Sao Miguel and Pico SM-7 Alkaline SM-7_rp1 Alkaline SM-7_rp2 Alkaline SM-7_rp3 Alkaline SM-7_rp4 Alkaline SM-7_rp5 Alkaline SM-27 Alkaline SM-31 Alkaline SM-33 Alkaline SM-36 Alkaline PI-6 Alkaline PI-10 Alkaline PI-12 Alkaline Average Sao Miguel and Pico σ Cape Verde FG-6 Alkaline FG-6_rp Alkaline FG-12 Alkaline FG-12_rp Alkaline SA-1 Alkaline SA-6 Alkaline SA-8 Alkaline SA-11 Alkaline Average Cape Verde σ La Palma LP-30a Alkaline LP-30a_rp Alkaline LP-113 Alkaline LP-45d Alkaline LP-107 Alkaline LP-111 Alkaline LP-134 Alkaline Average La Palma σ Hawaii ML-2-50 Alkaline LO Alkaline LO Alkaline LO Alkaline H 11 Alkaline Average Hawaii σ Continued. 21

22 Supplementary Information Table S3. ε 182 W and ε 183 W data for oceanic basalts and reference materials by single measurements (continued). Sample Chemistry 1 ε 182 W ±2σ m ε 183 W ±2σ m Ontong Java Plateau A-006R-06W Alkaline A-022R-03W Alkaline Average Ontong Java Plateau Manihiki-Plateau A-032R-01W Alkaline A-034R-01W Alkaline Average Shatsky Rise B-028R-02W Alkaline B-033R-05W_1 Alkaline B-033R-05W_2 Alkaline B-031R-03W Alkaline Average Shatsky Rise σ Iceland D-6 Alkaline Average oceanic basalts σ Reference materials BCR-2 BCR-2_1 Alkaline BCR-2_2 Alkaline BCR-2_3 Alkaline BCR-2_4 Alkaline BCR-2_5 Alkaline BCR-2_6 Alkaline Average BCR σ CPI W standard CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ Continued. 22

23 Supplementary Information Table S3. ε 182 W and ε 183 W data for oceanic basalts and reference materials by single measurements (continued). Sample Chemistry 1 ε 182 W ±2σ m ε 183 W ±2σ m CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ CPI_ Average CPI σ See method section 23

24 Table S4. Parameters used in hidden reservoir re-mixing model as compiled in refs. 25, 27, 29, Constant parameters Masses of reservoirs Continental crust (CC) Bulk silicate Earth (BSE) Value g g Isotope ratios and concentrations Nd system 147 Sm/ 144 Nd CHUR Nd/ 144 Nd CHUR Sm/ 144 T Sm CHUR ε 142 Nd EDR (relative to CHUR) [Nd] BSE 1.25 µg g -1 [Sm] BSE µg g -1 Isotope ratios and concentrations W system 182 Hf/ 180 T Hf CHUR W/ 184 W CHUR W/ 184 T W CHUR W/ 184 W BSE Hf/ 184 W CHUR 1.23 [Hf] BSE µg g -1 Atomic fraction a 180 Hf Atomic fraction a 184 W Atomic fraction a 182 W Decay constants λ 147 Sm a -1 λ 146 Sm a -1 λ 182 Hf a -1 Variable parameters Range Conc. Nd DM 0.5 to 1.0 µg g -1 Conc. Nd CC 18 to 22 µg g -1 ε 143 Nd DM 8 to 12 ε 143 Nd CC -13 to -20 ε 142 Nd DAE (relative to CHUR) 0.17 to 0.23 Conc. W BSE to µg g -1 X to

25 Table S5. Bulk mineral-melt distribution coefficients used in hidden reservoir re-mixing model for fractionation at upper and lower mantle conditions as compiled in ref. 4. Data taken from refs Sm Nd Hf W Olivine Orthopyroxene Clinopyroxene Garnet Mg-perovskite Bulk upper mantle Bulk lower mantle Olivine : Clinopyroxene : Orthopyroxene : Garnet = 0.57 : 0.29 : 0.00 : Mg-perovskite; value for W estimated using the strain lattice model published in ref

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