www.sciencemag.org/cgi/content/full/338/6109/939/DC1

Supplementary Material for

Evidence for a Dynamo in the Main Group Pallasite Parent Body

John A. Tarduno,* Rory D. Cottrell, Francis Nimmo, Julianna Hopkins, Julia Voronov, Austen Erickson, Eric Blackman, Edward R.D. Scott, Robert McKinley

*To whom correspondence should be addressed. E-mail: john.tarduno@rochester.edu

Published 16 November 2012, Science 338, 939 (2012)

DOI: 10.1126/science.1223932

This PDF file includes:

Materials and Methods

Figs. S1 to S7

Tables S1 to S5

References (33–65)

Tarduno et al., Evidence for a dynamo in the main group pallasite parent body

Supporting Online Material

Materials and Methods

Magnetic hysteresis data were collected using the University of Rochester Princeton Measure-

ments Corporation Alternating Gradient Force Magnetometer. Values for the examples shown in

Fig. 1 of the main text are as follows: Hcr, Hc and Mr/Ms are 154.6 Oe, 200.0 Oe and 0.3911

respectively for the Esquel specimen, and 111.1 Oe, 151.9 Oe and 0.3714, respectively, for the

Imilac specimen. For all remanence measurements we select mm-sized gem-like olivine subsam-

ples, lacking any surface discoloration that might be residual contamination from the surrounding

pallasite metal (we note that our initial tests revealed that samples with visible inclusions from

olivine crystal rims altered rapidly when heated). Obtaining suitable samples generally required

cleaning crystals in distilled water. A weak acid (HCl) was used on some crystals to remove surface

contamination. Remanence measurements were made with a 2G Enterprises 3-component 755R

DC SQUID magnetometer and a 2G small (6.3 mm) bore 3-component DC SQUID magnetometer

in the University of Rochester’s magnetically shielded room (ambient field <200 nT). CO2 laser

heating and cooling was conducted (in air) in additional magnetic shields to produce a magnetically

null environment.

Olivine samples 2-3 millimeters in size were mounted on the end of quartz tubes with Omega

cement (both of which are routinely measured to ensure the blank is in the 10−13 to 10−14 A m2

range). The sample holder also served as the target for CO2 laser heating (the 7 mm diameter laser

beam applied at peak temperature for ∼1 minute ensures uniform heating of the crystal; heatings

at each Thellier-Coe paleointensity step were for 3 minutes). The natural remanent magnetization

of approximately 15% of the clean crystal subsamples measured were in the 10−9 to 10−10 A m2

range; these are the focus of our studies as the magnetizations are well within the measuring

range of the DC SQUID magnetometers throughout the demagnetization procedures. The success

rate for crystals having these intensities (yielded interpretable paleointensity results) was ∼50%.

This compares well with paleointensity success rates from Thellier-Coe experiments on whole-rock

terrestrial basalts, which often average 20% (or less).

Thellier-Coe (24) paleointensity data consist of demagnetization of the NRM (field-off step),

followed by the reheating of the sample at the same temperature in a known applied field (field-on

step). We use orthogonal vector plots of the field-off steps to determine the optimal tempera-

ture range to calculate paleointensities. In this study, we typically use a lowermost Thellier-Coe

unblocking temperature for paleointensity calculation that is slightly higher than the lowest un-

blocking temperature where we believe a primary magnetization is held (i.e. 360 oC). This approach

is conservative, and aimed to avoid any influence of magnetizations held at lower unblocking tem-

peratures. For consistency, we use this same temperature range in determining paleointensity from

Total TRM data (see below), although we note that some minor alteration might be expected given

the cumulative time at elevated temperature.

1

Heatings were minimized by collecting Thellier-Coe paleointensity data only in the temperature

range where orthogonal vector plots show univectorial decay. An applied field of 60 µT was used for

all Thellier-Coe measurements. After NRM demagnetization and collection of Thellier-Coe data,

a Total TRM was applied. Using a CO2 laser, samples were heated to 700 oC and then cooled

in the presence of a field over a 10 minute time span. The Total TRM was subsequently stepwise

demagnetized using the CO2 laser. An applied field of 60 µT was used in the collection of all initial

Total TRM data. After demagnetization of the first Total TRM, subsample Imilac E3 was given

a second Total TRM in the presence of a 30 µT field (and subsequently demagnetized with a CO2

laser) to check for any potential applied field dependence on paleointensity.

To test for consistency in magnetic directions, an oriented section 1-mm thick was prepared.

Metal was etched away, leaving several mutually oriented gem-like olivine crystals, which we sub-

sequently separated (maintaining orientation) and thermally demagnetized using the CO2 laser.

SOM Text

Paleointensity selection criteria. Examples of accepted results are shown in Figure 2 of the main

text. Two additional examples of accepted results are included here (fig. S1). Results of Thellier-

Coe and Total TRM paleointensity experiments are reported in tables S1-2. Values are judged

acceptable if Thellier-Coe paleointensity and Total TRM paleofield estimates are consistent within

15% (see table S2). The uncertainty in the individual Thellier-Coe and Total TRM paleointensity

estimates must be ≤15%

Here we use demagnetization of a Total TRM to assess alteration because it can readily detect

(and in our case exclude) whole-scale transformations with heating seen in some FeNi magnetic

carriers in meteorites (10). Although our heating times using the CO2 laser are very rapid compared

to those of standard ovens used in paleomagnetism, we note that at the end of our experiments a

specimen has still been exposed to elevated temperatures for a cumulative time exceeding 2 hours.

We forgo pTRM checks (33) which, if applied, would have resulted in even longer cumulative times

at elevated temperature. The Total TRM data also aid in the interpretation of magnetizations

observed at high unblocking temperatures. For example, some Esquel olivine specimens acquire

additional partial TRMs after the temperature at which the NRM appears to have been completely

demagnetized. This is expressed as a flattening of NRM/TRM data (Fig. 2C), which in itself

might suggest that a very low (or null) field is recorded at high unblocking temperatures. However,

demagnetization of a Total TRM reveals only a minor TRM in this same temperature interval (Fig.

2D) suggesting that increases in partial TRM at high temperatures reflect either minor alteration

and/or the influence of minor, and more complex, magnetic phases (see discussion in “Minor high

unblocking temperature magnetizations” below).

Several factors contribute to the cause of unsuccessful experiments. The NRM intensity of

some samples decreased rapidly on AF demagnetization to levels after which measurement with

the SQUID magnetometers through an entire paleointensity run was no longer viable. The main

cause of unsuccessful samples that did not display such AF demagnetization characteristics appears

2

to be thermally-induced alteration. This was manifested by either a scattered NRM demagnetiza-

tion pattern (fig S2A,B) and/or a Total TRM curve that differed markedly from that of the NRM

demagnetization (fig S2B,C).

A B C

D

Inte

nsity x

10

-11 A

m2

Imilac

0

10

20

30

40

50

60

70

80

90

100

0 100 200 300 400 500 600 700 800

Total TRM

NRM

Esquel

410

Inte

nsity x

10

-11 A

m2

Temperature oC

360

410W E

N,Up

S,Down

0

5

10

15

0 5 10

500

410

NR

M (

x 1

0-1

1 A

m2)

TRM (x 10-11 A m2)

109.6 μT

0

20

40

60

80

100

350 400 450 500

Inte

nsity x

10

-11 A

m2

Temperature oC

113.4 μT

0

5

10

0 200 400 600 800

Temperature oC

400 500

NR

M (

x 1

0-1

1 A

m2)

TRM (x 10-11 A m2)

57.9 μT

400

500

0

1

0 1

E F G

H

W E

320

N,Up

S, Down

250

400

Temperature oC

Inte

nsity x

10

-11 A

m2

350 400 450 500

59.9 μT

0

1

2

3

Fig. S1. Additional examples of successful paleointensity experiments on pallasite olivine. (A)

Demagnetization of natural remanent magnetization (NRM) of Esquel olivine (black line). (B)

Orthogonal vector plot of (A), red is inclination, blue is declination (orientation relative). (C)

Thellier-Coe paleointensity data, NRM removed versus thermoremanent magnetization (TRM)

gained using a 60 µT applied field suggests a paleofield of 109.6 µT. (D) Demagnetization of a

laboratory Total TRM acquired in a 60 µT field (red curve in (A)) suggests a paleofield of 113.4

µT (calculated by comparing values at three temperature steps highlighted by grey boxes). (E-H)

Paleointensity data as discussed above on Imilac olivine indicating paleofields of 57.9 µT (Thellier-

Coe technique, 60 µT applied field) and 59.9 µT (Total TRM method, 60 µT applied field).

3

A B

C

0

2

4

6

8

10

12

14

N

S

W E

Temperature oC

Inte

nsity (

x10

−11 A

m2)

0 100 200 300 400 500 600 700

N

S

EW

0

10

20

30

40

50

60

70

80

Inte

nsity (

x10

−11 A

m2)

0 100 200 300 400 500 600 700 800

Temperature oC

0

2

4

6

8

10

12

14

Inte

nsity (

x10

−11 A

m2)

Temperature oC

0 100 200 300 400 500 600 700

N

S

W E

Fig. S2. Examples of paleointensity results that did not meet selection criteria. Intensity versus

temperature plots show natural remanent magnetization (NRM) decay (A,B,C) (black) and Total

Thermoremanent Magnetization decay (B,C) (red). Orthogonal vector plots are shown for NRM

demagnetization (red is inclination, blue is declination of relative orientation).

Paleointensity results and averages. Two pallasite meteorites were sampled (Esquel and Imi-

lac). Two thin slabs from each pallasite were available for study (denoted by 1, 2, respectively in

the tables below). Several consistency tests were performed and the results of these tests were in-

corporated into hierarchial averages (tables S3-4) as follows. For Total TRM, paleofield results from

the same crystal measured at different applied field values were averaged (“applied field average”).

Paleofield results from different subsamples from a single olivine crystal were averaged (“crystal

average”). Results from different crystals from a given meteorite sample were averaged (“meteorite

sample estimate”). “Meteorite averages” were determined by averaging the two meteorite sample

estimates available for each meteorite studied.

4

Table S1. Thellier-Coe paleointensity estimates.

Subsample FThC (µT) T (oC) [N] R2 f g

Esquel 1 (green4) 132.4 ±5.7 400-500 [3] 0.92 0.155 0.278

Esquel 2 (19c) 110.7 ±5.2 400-450 [3] 0.98 0.129 0.493

Esquel 2 (3c) 116.0 ±5.4 410-485 [6] 0.98 0.185 0.788

Esquel 2 (4b) 109.6 ±7.0 410-500 [7] 0.99 0.184 0.724

Imilac 1 (F8) 74.4 ±6.7 400-500 [3] 0.92 0.058 0.392

Imilac 1 (E3) 64.9 ±4.5 400-500 [5] 0.98 0.059 0.742

Imilac 1 (E7) 57.9 ±7.8 400-500 [5] 0.97 0.031 0.707

Imilac 2 (G9)* 82.1 ±6.3 425-520 [5] 0.98 0.038 0.739

Imilac 2 (G12) 79.3 ±7.2 400-520 [6] 0.94 0.081 0.726

Abbreviations: FThC , Thellier-Coe field value with 1σ uncertainty; T , temperature range of fit; N ,

number of temperature steps used in fit; f, g are fraction of NRM fit and gap factor, respectively,

from (33). ∗ Sample omitted from averages because of high Total TRM paleointensity uncertainty

(see table S2).

Table S2. Total TRM paleointensity estimates.

Subsample FTTRM (µT) T (oC) [N] ∆FTTRM−FThC%

Esquel 1 (green4) 134.3 ±6.1 400-500 [3] 1

Esquel 2 (19c) 118.8 ±5.7 400-450 [3] 7

Esquel 2 (3c) 115.9 ±6.8 410-485 [3] <-1

Esquel 2 (4b) 113.4 ±4.0 410-500 [3] 3

Imilac 1 (F8) 72.1 ±1.0 400-500 [3] -3

Imilac 1 (E3)† 65.9 ±4.4 400-500 [3] 2

Imilac 1 (E3)‡ 67.3 ±3.4 400-500 [3] 4

Imilac 1 (E7) 59.9 ±1.0 400-500 [3] 3

Imilac 2 (G9)* 84.5 ±15.5 425-520 [3] 3

Imilac 2 (G12) 77.7 ±2.2 400-520 [3] - 2

Abbreviations: FTTRM , Total TRM field value estimate with 1σ uncertainty; T , temperature range

of fit; N , number of temperature steps used in fit; ∆FTTRM−FThC, difference between Total TRM

and Thellier-Coe paleointensity estimates, expressed as percent of the Thellier-Coe value. † 60 µT

applied field; ‡ 30 µT applied field. ∗ Sample omitted from averages because of high Total TRM

paleointensity uncertainty.

5

Table S3. Thellier-Coe hierarchical paleointensity averages.

Subsample FThC (µT) Crystal Meteorite sample estimate Meteorite

average (µT) (µT) average (µT)

Esquel 1

green4 132.4 ±5.7 132.4 Esquel

Esquel 2 122.3 ±14.4

19c 110.7 ±5.2 112.1 ±3.4 (N=2)

3c 116.0 ±5.4 (N=3)

4b 109.6 ±7.0

Imilac 1

F8 74.4 ±6.7 67.9 ±9.2 Imilac

E3 64.9 ±4.5 61.4 ±4.9 (N=2) 73.6 ±8.1

E7 57.9 ±7.8 (N=2) (N=2)

Imilac 2

G12 79.3 ±7.2 79.3

Abbreviations: FThC , Thellier-Coe field value. All averages shown with 1σ uncertainty.

Table S4. Total TRM hierarchical paleointensity averages.

Subsample FTTRM (µT) Applied field Crystal Meteorite sample Meteorite

average (µT) average (µT) estimate (µT) average (µT)

Esquel 1

green4 134.3 ±6.1 134.3 Esquel

Esquel 2 125.2 ±12.9

19c 118.8 ±5.7 116.0 ±2.7 (N=2)

3c 115.9 ±6.8 (N=3)

4b 113.4 ±4.0

Imilac 1

F8 72.1 ±1.0 67.7 ±6.2

E3† 65.9 ±4.4 66.6 ±1.0 (N=2)

E3‡ 67.3 ±3.4 (N=2) 63.3 ±4.7

E7 59.9 ±1.0 (N=2) Imilac

Imilac 2 72.7 ±7.1

G12 77.7 ±2.2 77.7 (N=2)

Abbreviations: FTTRM , Total TRM field value estimate. All averages shown with 1σ uncertainty.† 60 µT applied field; ‡ 30 µT applied field.

6

Minor high unblocking temperature magnetizations. Although the dominant natural re-

manent magnetization is removed by thermal demagnetization between 360 and 500 oC, consistent

with a taenite carrier, we note there is a very small signal (1-5% of the NRM) at demagnetization

temperatures >500 oC in some samples. On the basis of microprobe analyses (discussed below) and

potential unblocking temperatures, we consider these small signals to be carried by a fine-grained

mixture of taenite and kamacite. We further note that some samples show a small NRM and To-

tal TRM remanence increase (and subsequent decrease) at thermal demagnetization temperatures

>500 oC (cf Figure 2). This increase generally occurs over a restricted temperature range (∼100oC), but its exact initiation temperature varies between samples. We interpret this as reflecting ex-

change interaction between fine-grained taenite and kamacite. Because these are very minor phases

compared to the bulk magnetization, this interaction is not apparent in FORC diagrams. We also

note that small amounts of tetrataenite could be recorded at these high unblocking temperatures.

However, the reproducibility of the intensity increase seen in demagnetization of a Total TRM (see

Figure 2e) indicates that tetrataenite cannot be solely responsible for these minor magnetizations

because tetrataenite should not have survived heating to 700 oC (i.e. the temperature at which the

Total TRM was applied).

Terrestrial weathering. Unblocking temperatures similar (but not identical) to those reported

in our study have been reported by Uehara et al. (34) in weathered chondrite meteorites and

interpreted to reflect maghemite and substituted magnetite formed during terrestrial weathering,

resulting in a terrestrial magnetization overprinting an extraterrestrial signal. This was not the

case for chondrites with no or little weathering. Maghemite generally inverts after heating above

250 oC (21), and this results in irreversible magnetic behavior; this was not observed in our thermal

demagnetization experiments. Moreover, evidence for maghemite or a substituted magnetite phase

was not found during our SEM or microprobe analyses (detailed below), whereas clear evidence

for FeNi particles was identified. However, we emphasize that our analyses have been restricted to

gem-like olivine particles. Our meteorite samples were selected to have minimal weathering. Al-

though not studied here, we predict that weathered pallasite olivines do contain magnetic minerals

formed during terrestrial weathering.

SEM and Microprobe analyses of FeNi particles. Scanning electron microscopy (SEM)

analyses were conducted using a Zeiss SUPRA 40VP with EDAX spectrometer at the University

of Rochester. SEM analyses reveal FeNi inclusions that are potential remanence recorders. These

are similar to those reported in some prior studies (35-36) but differ from the tubular symplectic

inclusions studied in the Fukang pallasite (37). We observed some Cr-rich inclusions, but these

are not candidates for the major NRM carrier which demagnetizes between 360 and 500 oC. SEM

analyses of an olivine inclusion that is a candidate remanence carrier from the Esquel meteorite is

shown in fig. S3.

7

Maps of inclusion 7 in crystal D2

FeK FeL

Mg

NiL NiK

OSi

S

C

Esquel - Crystal D2, Inclusion 7

Fig. S3. SEM analyses of an inclusion in olivine of the Esquel pallasite meteorite. EDAX K and

L shell shell composition maps are shown for Fe and Ni.

EDAX spectra show an absence of Si, Mg and O, indicating that the inclusion is distinct from the

olivine matrix. Sulfur-rich regions (darker grey areas of the inclusion in the SEM image) separate

concentrations of FeNi within the inclusion.

Compositions of inclusions were further explored using a JEOL 8900 electron microprobe at

Cornell University with an accelerating voltage of 8 KeV to obtain ∼0.5 micron resolution. Electron

microprobe results reveal FeNi compositions within the inclusion (fig. S4). A pentlandite (Fe,

Ni)9S8 standard from Manibridge, Ontario (weight percentages S: 33.01, Fe: 30.77, Co: 0.10, Ni:

36.12) was used for these analyses. Total weight percentages less than 100% in the analyses plotted

reflect the presence of elements other than Fe and Ni (mostly S). The compositions of Ni-rich

particles overlap with those of the ordered FeNi mineral tetrataenite. However, the dominant

8

changes in NRM intensity do not match the characteristic magnetic decay pattern related to the

∼550 oC Curie temperature of tetrataenite (38) (see also fig. S5). In addition, as noted above, the

reproducibility exhibited by the Total TRM data (Figure 2, tables S2, S4) are inconsistent with

tetrataenite.

Esquel - Crystal D2 Inclusion 7 - Microprobe Analysis

2µm

A

D

F

B

E

G

C

H

0

1000

2000

3000

4000

5000

6000

7000

8000

9000

10000

11000

12000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Fe

C

S FeNi AlMg

0

1000

2000

3000

4000

5000

6000

7000

8000

9000

10000

11000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Ni

Fe

C

SAlFeSi

O

0

1000

2000

3000

4000

5000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Fe

C

S FeNi

0

1000

2000

3000

4000

5000

6000

7000

8000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Fe

C

S

MgFe

S

S

0

1000

2000

3000

4000

5000

6000

7000

8000

9000

10000

11000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Ni

Fe

C

SAl FeMg Si

0

1000

2000

3000

4000

5000

6000

7000

8000

9000

10000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Ni

Fe

C

SAl FeMg Si

Ni

Fe

CS

Al FeO

0

1000

2000

3000

4000

5000

6000

7000

8000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Fe

C

S Fe

Ni

AlO

Ca0

1000

2000

3000

4000

5000

6000

7000

8000

9000

10000

11000

12000

Coun

ts

0 1 2 3 4 5 6 7 8

keV

Ni Fe Total Weight 57.022 42.845 99.867Atomic 55.8707 44.1293 100

Ni Fe Total Weight 58.136 42.209 100.345Atomic 56.7148 43.2852 100

Ni Fe Total Weight 52.25 44.877 97.127Atomic 52.5523 47.4477 100

Ni Fe Total Weight 3.237 101.377 104.614Atomic 2.9478 97.0522 100

Ni Fe Total Weight 3.015 99.185 102.2Atomic 2.8103 97.1897 100

Ni Fe Total Weight 52.702 41.762 94.464Atomic 54.5557 45.4443 100

Ni Fe Total Weight 0.126 45.685 45.811Atomic 0.2622 99.7378 100

Fig. S4. Electron microprobe results for Esquel olivine (same inclusion shown in Fig. S4).

Brightest areas (B,F,G,H) are high Ni, surrounded by Fe sulfide (A). Slightly less bright areas are

low Ni, FeNi metal (C,D). Area (E) likely samples region of several adjacent grains, but nonetheless

demonstrates S between region of low and high Ni.

9

0

20

40

60

80

0 200 400 600 800

Inte

nsity

x10

-10

A m

2

Temperature oC

Observed NRM

Observed Total TRM

400

400

Predicted tetrataenitedemagnetization

Fig. S5. Predicted tetrataenite NRM demagnetization behavior versus observed demagnetization

of NRM and total TRM of an Esquel pallasite olivine subsample (cf. Figure 2). Note the NRM

value of the hypothetical tetrataenite component is arbitrary, and the decay qualitative, drawn

after the behavior reported by Wasilewski (38).

We interpret the dominant magnetic recorder in inclusions within the pallasite olivine as

metastable, disordered taenite. We note that this conclusion differs from the result of a thermo-

magnetic analysis of the “silicate” (presumable olivine) part of a pallasite (Yamato-74044) reported

in an early study by Nagata (11) which is consistent with kamacite. We note, however, that our

analyses have been restricted to gem-like olivines lacking large inclusions. It is possible that con-

tamination from the metal surrounding the olivine (especially near the olivine rim) was responsible

for the signal reported by Nagata (11).

Given the slow cooling recorded by the massive FeNi surrounding the olivine (28), we might ex-

pect taenite to have ordered. However, the high-Ni magnetic carriers in olivine have a very different

origin from the tetrataenite rims in the metallic matrix as the latter formed by low temperature

equilibration between FeNi phases whereas the former are minor components in FeS inclusions that

are predominantly composed of FeS. We envisage that the Fe-Ni-S inclusions were injected as a

melt into cracks [perhaps at ∼900 oC, (39)] which were sealed up on cooling. We speculate that

minor Fe exchange between the FeNi and FeS phases during slow cooling may inhibit ordering in

very small particles such as those that characterize the olivine magnetic inclusions.

Pressure effects on magnetization. Hydrostatic pressure can potentially affect magnetizations

(21). Turcotte and Schubert [(40), eq. 2-71] give the pressure as a function of radial position r

inside a homogeneous body of density ρ:

P =2

3πρ2G(R2 − r2) (1)

where R is the radius of the body and G is the gravitational constant. Taking our estimates of the

size of the pallasite parent body, R=200 km, ρ=3 g/cc, we obtain pressures of 9.6 MPa at r =180 km

10

and 22 MPa at r=150 km. Although detailed pressure experiments are not available on sub-micron

sized FeNi particles at these estimated pressures, we note that our estimates of pressures are much

lower than the value at which substantial changes are observed in magnetite and titanomagnetite

[e.g. (41-42)] and FeNi (43). Shock can induce much larger pressures. However, beyond the event

that caused olivine-metal mixing and break-up of the parent body, pallasites do not appear to

have been affected by the massive shocks that characterize other meteorites. For example, pallasite

metal does not show the characteristic hatched pattern seen in shocked IIIAB iron meteorites (44 ),

and dislocation densities are low in both olivine and metal [e.g. (9, 45)].

Martelli and Newton (46) studied shock associated with a hypervelocity projectile (15 km/s)

cratering a terrestrial basalt and proposed that a shock magnetization comparable to a TRM could

be produced. In this case, the shock magnetization was thought to be related to a plasma formed

by the impact. Because the pallasites we have studied lack massive shock features, we believe they

were not near the surface at the time of the impact that disrupted the parent body. In this case, the

more relevant far-field remagnetization process is shock-induced crystallographic transformation.

Dickinson and Wasilewski (47) found that this process normally induces a magnetization much less

than that recorded by our TRM experiments.

Absolute ages of metal-olivine mixing. Lugmair and Shukolyukov (26) report 53Mn-53Cr

systematics on main group pallasites Omolon and Springwater, suggesting ages of 4558.0 ± 1.0 Ma

and ∼4557 Ma, respectively. These ages are consistent with those reported from Re-Os systematics

by Shen et el. (48) (4.60± 0.5 Ga), who interpret the time scale for pallasite fractional crystalization

as spanning 10-20 Myr. These ages provide an upper bound (i.e. oldest age) on the time of metal-

olivine mixing.

In an early K-Ar study, Megrue (49) reported ages of 4.3 Ga for 3 pallasites (Krasnojarsk,

Marjalahti, and Springwater). A much younger 40Ar-39Ar age of 0.86 Ga has been reported for

the non-main group pallasite Eagle Station by Niemeyer (50), and interpreted as recording an

impact of that age. But that study was conducted on weathered olivine [(50), p. 1007] and the age

probably reflects terrestrial processes. Similarly Megrue (49), noted that an apparent young Kr-Ar

age of 0.44 Ga for the Admire pallasite could reflect “diffusion of radiogenic Ar40” or “potassium

contamination of the sample by terrestrial weathering”.

Bondar and Perelygin (27) reported model fission track ages of 4.37 Ga ± 0.01 Ga for the

Marjalahti pallasite, 4.19 ± 0.02 Ga for the Omolon pallasite, 4.18 ± 0.03 Ga for the Bragin pall-

asite and 4.21 ± 0.02 for the Krasnoyarsk pallasite. Given that the temperature for the complete

removal of tracks is 450-500 oC (27), these ages could be close to the time of acquisition of the

magnetizations we have recorded. However, such an assignment depends critically on the accuracy

of the fission track age model. Nevertheless, these ages provide a lower bound (i.e. youngest age)

on the time of metal-olivine mixing.

Thermal modeling. Given the uncertainties in even such basic parameters as its size, we adopt

11

a very simple model to track the thermal evolution of the pallasite parent body. This model re-

sembles those of previous investigations [e.g. (29, 51-53)]; one important complication, however,

is that it includes an insulating, near-surface megaregolith (54-55). Although, especially for the

larger bodies, mantle convection or advection via melt might be important in the early stages, we

simply assume spherically-symmetric conductive transfer of heat:

∂T

∂t=

1

r2∂

∂r

(κr2

∂T

∂r

)(2)

where T (r) is temperature, r is radial position, and κ is thermal diffusivity. We neglect internal

heat production because of the ∼100 Myr timescale of greatest interest to us. Since this timescale

is long compared to the decay of 26Al or 60Fe, the effect of this early heating is incorporated into

our initial conditions. Conversely, over the timescale of interest the decay of long-lived radiogenic

elements like 40K will not qualitatively change our conclusions.

The temperature at the surface (r = R) is kept at a fixed value Ts. The temperature gradient

at r = 0 is zero. We assume an isothermal initial condition (T=1600 K) in which the mantle is

solid but the core is liquid. We solve equation (2) numerically using a centered finite-difference

scheme.

The near-surface megaregolith layer is assumed to have a thermal diffusivity one-tenth that of

the solid rock (54). Based on Fig. 12 of Warren (55), we assume megaregolith thicknesses of 5, 8

and 24 km for bodies of radii 100, 200 and 600 km, respectively.

The megaregolith is important not only as an insulator, but also because it is an unlikely source

of the pallasites (which would have experienced shock and mixing with other silicate phases, neither

of which are observed). We consider large (600 km) objects as unlikely sources of pallasites because

the cooling rate data would require them to have originated within the megaregolith layer (see main

text). We also note that it is possible to ultimately disrupt the parent body by impact without

widespread heating (56).

If core solidification has not yet finished, we only solve equation (2) within the mantle and take

the temperature at the core-mantle boundary (r = Rc) as our boundary condition. The core is

assumed isothermal (due to vigorous convection) and its bulk temperature is updated based on the

heat extracted into the mantle (4πR2ckm

∂T∂r |r=Rc ∆t ) during that timestep ∆t , where km is the

mantle thermal conductivity. We take Rc = R/2.

We assume that the core solidifies at a single temperature Tl, here taken to be 1200 K. If T = Tl

during a particular timestep, the core temperature then remains pinned at that temperature until

the total latent heat of the core (43πR3cρcLc) has been extracted into the mantle via conduction.

Here Lc is the specific latent heat of the core material and ρc its density.

Once the core has completely solidified, we then proceed to solve equation (2) as before, but

now the bottom boundary condition is to set the temperature gradient to zero at r=0. Note that

the thermal diffusivities κm and κc for mantle and core respectively are quite different.

We use a grid spacing ∆r =R/100 and adopt a constant timestep governed by the Courant

criterion: ∆t =0.3∆r2/κc, where we use the core diffusivity because it is more restrictive. We

12

verified that our numerical model reproduces the analytical solutions for a uniform sphere cooling

from an isothermal initial state (57) when the core radius is zero. Parameter values adopted are

given in table S5.

This model ignores many details, such as the fact that core solidification is likely a multi-phase

process which involves a progressive reduction in solidus temperature as solidification proceeds.

Nonetheless, the conclusions reached in the main text are unlikely to be significantly affected by

such uncertainties. This is because the key results - cooling rate and liquid core lifetime - are

both primarily dependent on the size of the parent body, while a low blocking temperature implies

relatively shallow depths, irrespective of the details of the calculations.

In the main text we focus on the 2-9 K/Myr cooling data for main group pallasites obtained by

Yang et al. (28). We do not consider one higher rate reported in that study (18 K/Myr ± 9 K/Myr)

because of its high uncertainty; it also appears to be an outlier in cooling rate, and in cloudy zone

particle size versus tetrataenite bandwidth (Fig. 7 of Yang et al., (28)). Ito and Ganguly (58)

calculate a cooling rate of 20-40 K/Myr at 1000◦C on the basis of a thermochronological model

applied to 53Mn-53Cr age data on olivine from the Omolon pallasite (26). For our 200 km radius

model object, this cooling rate is obtained over a depth range of 6-20 km, which is compatible with

the results shown in Fig. 3 of the main text.

Because we envisage the metallic injections as being relatively narrow, dike-like features, they

will cool rapidly to attain the background temperature of the surrounding material. Their subse-

quent cooling history (including passage through the Curie temperature) will then simply be that

of the surrounding material.

We also considered the possible role of impact heating during the injection event, but concluded

that it was likely to have negligible effect on the overall cooling, as follows. Early in the history of the

asteroid belt, impact velocities were probably comparable to escape velocities. The escape velocity

is roughly 200 m/s (R/100 km) where R is the radius of the body. So for bodies a few hundred

km in radius, impact velocities will be lower than the sound speed of intact rock (a few km/s), i.e.

the impacts are subsonic. As a result, the impact heating (and associated shock pressures) will be

small. If all the kinetic energy of a 200 m/s projectile were contained within one projectile radius,

the temperature increase would only be about 20 K and the actual temperature change would be

even less, because the heat is in reality more broadly distributed.

Table S5. Thermal modeling parameters, based on Table 1 of Sahijpal et al. (52).

Quantity Value Units Quantity Value Units

Ts 250 K Tl 1213 K

ρm 3000 kg m−3 ρc 7800 kg m−3

κm 5× 10−7 m2 s−1 κc 5× 10−6 m2 s−1

Lc 0.27 MJ kg−1 km 3 Wm−1K−1

13

Late impact heating and remagnetization. Many HED meteorites have 39Ar-40Ar ages that

suggest impact-induced heating (59-60) to temperatures capable of partially (or completely) reset-

ting a magnetization that might have been acquired during initial cooling. Some of these impact

heating events occur as late as ∼3.4 Ga, at a time when dynamo action might have ceased on the

HED parent body (Vesta) due to core solidification. If the pallasite meteorites we have studied

were reheated in a similar manner, after dynamo cessation on the pallasite parent body, it would

affect our conclusions. Here, we first consider how our conclusions would be affected, and then

explore whether impact reheating is consistent with available constraints.

If the pallasites were reheated after dynamo action ceased, an ambient magnetic field still must

have been present to account for the magnetizations we have recorded. A potential source of this

ambient field is the entire parent body mantle and crust, which if magnetized during the dynamo

epoch would create a global remanent field. The relevant magnetic carrier of the magnetization

would be kamacite, since its high Curie point (approximately 765 oC, (21)) is consistent with the

idea that it could retain a magnetization even if heated to temperatures sufficient to reset the

magnetizations we have reported from the pallasites.

In this case, the constraints on the pallasite parent body would come from the cooling rates of

the pallasite metal and the inferred magnetization in the parent body mantle and crust (i.e. not

directly from the magnetization of the pallasite olivine, which in this scenario is reset by impact

heating and records only an echo of the dynamo). Given these new constraints, a 600 km radius

body can still be excluded because the pallasites would have to reside in the regolith. However,

we would have less resolution on the position of the pallasites within the parent body and on the

parent body size. Specifically, the pallasites could have formed deeper in the parent body, within

10% of the core-mantle boundary, and parent bodies ranging in size from 100 to 200 km radius

could satisfy the constraints.

Paleomagnetic directional and paleointensity data from the pallasite olivine are linear between

360 oC and 500 oC (Fig 2). It is unlikely that the ambient magnetic field during a hypothetical late

heating event induced by impact, and the field during initial cooling would be exactly the same.

Hence, following Thelliers’ laws of pTRM additivity (21), if the hypothetical late heating event

reached temperatures greater than 360 oC, but less than 500 oC, a break in slope would be seen in

the paleointensity and/or directional data. Because this is not observed, we conclude that any late

stage heating must have reached a temperature close to 500 oC to be compatible with the magnetic

data.

As noted above, a low velocity impact is insufficient to generate a temperature change great

enough to affect the magnetization. High-velocity impacts create transient high pressures; as the

material subsequently unloads along an adiabat, heating results with higher peak pressures causing

more heating (61). To quantify this effect we used the same approach and parameters as in Barnhart

et al. (62), Appendix A2. We calculate the peak pressure and temperature increase ∆T experienced

directly below the impactor. The results are plotted in Fig S6. As expected, there is an almost

linear relationship between peak pressure and temperature increase. A temperature increase to 500

14

oC is required to reset the magnetization. This implies temperature increases ranging from 300 K

for an impact shortly after core solidification to greater than 400 K for impacts >1 billion years

later. This heating range implies shock pressures ranging from 4.5 GPa to greater than 5.5 GPa

(∼45-55 kbar).

Studies of olivine dislocation density in main group pallasites limit shock pressure to a few

tens of bars (9). These shock levels are incompatible with the much higher shock (i.e. GPa’s)

that would accompany impacts needed to produce heating sufficient to reset the magnetizations we

have recorded. These studies were conducted only on the Admire, Brenham and Dora pallasites.

But metal macro- and microstructure is also sensitive to shock and shock-induced heating. For

example, shock up to 1 GPa can cause Neumann band (mechanical twins) in kamacite, and cloudy

taenite intergrowth can be completely lost upon short term heating (years) at 500 oC (63). Metal

microstructure has been examined in detail for the Esquel and Imilac pallasites (28) and no evidence

for shock has been reported.

Because relatively high shock accompanies impact heating (Fig. S6), and there is no evidence

of shock at these levels in the pallasites following the olivine-metal mixing event, we conclude that

resetting of the pallasite magnetization by impacts is unlikely. We therefore predict that any future

Ar-Ar or fission track dating efforts on the Esquel and Imilac pallasites should yield ages at least

as old as the 4.4-4.2 Ga ages (27) obtained from other main group pallasites.

Figure S6. Shock pressure and temperature increase as a function of impact velocity vi, calcu-

lated using the approach of Barnhart et al. ((62), Appendix A2). We assumed target and projectile

densities of 2.6 g/cc. Other parameters are as given in Barnhart et al. (62).

15

Model for parent body evolution. Our model for the origin of pallasites and the nature and

evolution of their parent body is based on paleomagnetic constraints, cooling rate data (28), thermal

modeling, and previously published geochemical analyses. We envision the impact of a differen-

tiated body with a partially solidified core into a larger differentiated body having a liquid core.

Liquid FeNi from the impactor intrudes into target dunite mantle in the form of dikes, which result

in the formation of the characteristic pallasite compositions. This source of FeNi metal satisfies

Ir constraints (3). The mantle and crust of the impactor are lost and/or incorporated into the

shallowest levels of the target. A similar impact scenario for the generalized case of planetesimal

accretion is presented by Asphaug et al. [(64) Figures 6-7, panels d-f of that paper]. It is also pos-

sible that the impactor had a very thin mantle/core, especially if it had seen prior impacts. There

are more complex scenarios that might meet the available constraints, but our central conclusions

are that the main group pallasites cooled in a body with a core dynamo and that the metal in

the pallasites did not come from that core. Our preferred model (fig. S7) is arguably the simplest

explanation of these findings.

Fig. S7 (next page). Our model for the origin of pallasites and the nature and evolution

of their parent body is based on paleomagnetic constraints and thermal modeling. (A) Liquid

FeNi from an impactor is injected into the mantle of the parent body. (B) Enlargement of boxed

region in (A) shows injection of metal in dike-like intrusions in dunite mantle, which result in the

pallasite texture (C) During the dike intrusion, metal is injected into fractures in the olivine (D);

these fractures subsequently heal upon cooling. Although we illustrate one impact, several similar

impacts might have occurred, and they together might explain the scatter in Ir values seen in main

group pallasites [e.g. (65)]. (E) Heat flux and temperature versus time. At a reference of 30 km

depth, the lower range of taenite Curie temperatures (estimated from the observed paleomagnetic

unblocking temperatures) will be reached at ∼137 million years after initial cooling. Core heat

flux is sufficient at this time to drive a dynamo if compositional convection occurs (14). (F) After

cooling below the lower Curie temperatures of taenite, collision of the pallasite parent body with

a similarly-sized protoplanet occurred, possibly in the terrestrial planet-forming zone [e.g. (32)],

destroying the pallasite parent body. (G) Pallasite meteorites were later scattered by interaction

with a protoplanet.

16

mag

netic

field

magnetic

field

0

50

100

T

ime,

Myr

150

200

250

300

2

0 2

00

4

0 4

00

6

0 6

00

8

0

coo

lin

g r

ate

reco

rded

b

lock

ing

tem

per

atu

re

8

00

-2 Heat flux out of core, mWm

100

C

ore

hea

t fl

ux

core

so

lid

ific

atio

n

star

ts

core

so

lid

ific

atio

n

en

ds

T

emp

erat

ure

(3

0 k

m d

epth

) 1000

120

Temperature, K

1200

~1

00

km

~1

cm

~1

0 μ

m

FG

BA

C

E

Fig

S7

D

17

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