# Carbonate record of temporal change in oxygen fugacity and gaseous species in asteroid Ryugu

Wataru Fujiya, Noriyuki Kawasaki, Kazuhide Nagashima, Naoya Sakamoto, Conel M. O'D. Alexander, Noriko T. Kita, Kouki Kitajima, Yoshinari Abe, Jérôme Aléon, Sachiko Amari, et al.

► To cite this version:

Wataru Fujiya, Noriyuki Kawasaki, Kazuhide Nagashima, Naoya Sakamoto, Conel M. O'D. Alexander, et al.. Carbonate record of temporal change in oxygen fugacity and gaseous species in asteroid Ryugu. Nature Geoscience, 2023, 16, pp.675-682. 10.1038/s41561-023-01226-y. insu-04187886

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# Carbonate record of temporal change in oxygen fugacity and gaseous species in asteroid Ryugu

Received: 21 December 2022

Accepted: 14 June 2023

Published online: 10 July 2023

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A list of authors and their affiliations appears at the end of the paper

The Hayabusa2 spacecraft explored asteroid Ryugu and brought its surface materials to Earth. Ryugu samples resemble Ivuna-type (CI) chondrites— the most chemically primitive meteorites—and contain secondary phyllosilicates and carbonates, which are indicative of aqueous alteration. Understanding the conditions (such as temperature, redox state and fluid composition) during aqueous alteration is crucial to elucidating how Ryugu evolved to its present state, but little is known about the temporal changes in these conditions. Here we show that calcium carbonate (calcite) grains in Ryugu and Ivuna samples have variable ¹⁸O/¹⁶O and ¹³C/¹²C ratios that are, respectively, 24–46‰ and 65–108‰ greater than terrestrial standard values, whereas those of calcium–magnesium carbonate (dolomite) grains are much more homogeneous, ranging within 31–36‰ for oxygen and 67–75‰ for carbon. We infer that the calcite precipitated first over a wide range of temperatures and oxygen partial pressures, and that the proportion of gaseous CO₂/CO/CH₄ molecules changed temporally. By contrast, the dolomite formed later in a more oxygen-rich and thus CO₂-dominated environment when the system was approaching equilibrium. The characteristic isotopic compositions of secondary carbonates in Ryugu and Ivuna are not observed for other hydrous meteorites, suggesting a unique evolutionary pathway for their parent asteroid(s).

The Japan Aerospace Exploration Agency Hayabusa2 spacecraft explored the near-Earth asteroid 162173 Ryugu and brought samples of its surface materials back to Earth¹,². Ryugu has been classified spectroscopically as a member of the C-complex asteroids³,⁴. It is a rubble-pile asteroid consisting of numerous rocky blocks that are the fragments resulting from the disruption of an original, larger parent body⁵⁻⁷.

Previous work has reported that Ryugu materials underwent extensive aqueous alteration as the result of water activity in the original parent body⁸⁻¹⁰ and are composed mainly of secondary minerals that formed during the aqueous alteration: phyllosilicates, carbonates, sulfides and oxides. Primary minerals such as anhydrous silicates are rare⁹. These petrological characteristics are comparable to those in the CI (Ivuna-type) chondritic meteorites, pointing to a kinship between

Ryugu and CI chondrites. The Ryugu's whole-rock chemical and isotope compositions confirm a close affinity with CI chondrites⁸,¹¹,¹².

Carbonates, the major Ca budget in Ryugu and CI chondrites¹³, are of particular interest because (1) their chemical and isotopic compositions reflect the conditions of aqueous alteration and (2) their grain size is commonly large enough to allow in-situ analysis by electron and ion microprobes¹⁴⁻¹⁷. The C source of carbonates is unclear but was probably ices that included CO, CO₂ and CH₄ and/or organic matter¹⁸,¹⁹. These materials may have formed in the solar nebula or even the parental molecular cloud of the Solar System. Therefore, the C isotope compositions recorded by carbonates can help us shed light on the physicochemical processes that operated in these environments.

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e-mail: wataru.fujiya.sci@vc.ibaraki.ac.jp
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<figure><img src="image_5.png" /><figcaption>**Fig. 1** | Calcite grains in the Ryugu C0002 sample. **a**, Mg (red), Si (green) and Ca (blue) elemental maps showing the distribution of calcite grains. Calcite grains (blue) in the Ryugu and Ivuna samples can be found in limited areas with Mg-rich olivine (yellow) and pyroxene (light green). Purple grains are dolomite. The large red grain is breunnerite. **b**, Backscattered electron image of a calcite grain.</figcaption></figure>

In this Article, we investigate the conditions of aqueous alteration and the origin of the materials accreted by the Ryugu/CI parent bodies. To this end, we performed in-situ O and C isotope measurements of calcite (CaCO₃) and dolomite (CaMg(CO₃)₂) in Ryugu samples A0058 (collected at the first touchdown site) and C0002 (from the second touchdown site) as well as the Ivuna meteorite. The O isotope data of the A0058 and Ivuna dolomite are taken from a previous study⁸.

## Occurrences and isotope compositions of carbonate minerals

Dolomite is the most abundant carbonate mineral in the analysed samples, and we found numerous dolomite grains throughout the Ryugu and Ivuna matrices, whereas calcite is rare. No calcite grains were observed in the Ryugu A0058 sample studied. Calcite was found only in limited areas (clasts) of the Ryugu C0002 and Ivuna samples, occurring with primary anhydrous silicate minerals such as Mg-rich olivine and pyroxene (Fig. 1a). The calcite grains (<10 µm in size) are usually smaller than the dolomite grains (several tens of micrometres) (Fig. 1b). The dolomite has compositional variation and complex zoning within grains (Extended Data Fig. 1). We also found breunnerite (Mg(Fe,Mn)(CO₃)₂) grains, for which we did not measure isotope compositions because of the lack of a suitable standard material for isotope analysis.

The O and C isotope compositions of the Ryugu and Ivuna carbonates are similar (Table 1). The δ¹⁸O values of the calcite in C0002 and Ivuna (+24‰ to +46‰) show a grain-to-grain variation larger than those of the dolomite in C0002, A0058 and Ivuna (+26‰ to +31‰) (δ'O in ‰ = [(('O/¹⁶O')sample/('O/¹⁶O')VSMOW - 1] × 1,000; i = 17 or 18, and VSMOW is the terrestrial standard material Vienna standard mean ocean water) (Fig. 2a). The simple average of the Δ¹⁷O values, the deviation from the terrestrial fractionation line defined as $\Delta^{17}\text{O} = \delta^{17}\text{O} - 0.52 \times \delta^{18}\text{O}$, of the C0002 and Ivuna calcite is +1.37 ± 0.40‰ (2 s.e., N = 17). The Δ¹⁷O values of the C0002, A0058 and Ivuna dolomite are systematically lower than those of the calcite, and the average Δ¹⁷O value of the dolomite (+0.26 ± 0.23‰, 2 s.e., N = 16) is closer to the whole-rock values of three Ryugu samples⁸ (+0.61 ± 0.28‰, 2 s.d.; Fig. 2b). The previously measured O isotope compositions of the A0058 and Ivuna dolomite are in good agreement with those of the C0002 dolomite measured in this study⁸.

Like the δ¹⁸O values, the dolomite in A0058, C0002 and Ivuna has a relatively narrow range of δ¹³C values, from +67‰ to +75‰ (δ¹³C in ‰ = [(¹³C/¹²C)sample/((¹³C/¹²C)VPDB - 1)] × 1,000; VPDB is the terrestrial standard material, Vienna PeeDee Belemnite) (Fig. 3). The δ¹³C and δ¹⁸O values of the dolomite are broadly consistent with the bulk δ¹³C and δ¹⁸O values of carbonates, including calcite, dolomite and breunnerite, in CI chondrites (Ivuna and Orgueil) determined on whole-rock samples¹⁸, corroborating the observation that dolomite is the major carbonate mineral. However, the δ¹³C values of the calcite in C0002 and Ivuna are highly heterogeneous from grain to grain, ranging from +65‰ to +108‰, and they are commonly higher than those of the Ryugu and Ivuna dolomite (Fig. 3). These variations have also been reported for other Ryugu samples²⁰, confirming that our results represent the isotopic characteristics of Ryugu carbonates. For calcite grains on which we conducted multiple measurements, the δ¹³C values in each grain are identical within uncertainties of -5.1‰. Thus, the heterogeneity of δ¹³C values within individual calcite grains is probably no larger than 10‰, which is much smaller than the heterogeneity of >40‰ between grains.

## Isotope signatures of the Ryugu and CI carbonates

If the carbonates were in O and C isotopic equilibrium with the aqueous fluid when they precipitated, their O and C isotope compositions would have been determined by the mass-dependent equilibrium isotopic fractionation between carbonates and water for O, and that between carbonates and dissolved CO₃²⁻ (and other dissolved C-bearing chemical species) for C. The magnitude of this equilibrium isotopic fractionation depends on temperature²¹. Thus, the O and C isotope compositions of the carbonates would reflect those of water and CO₃²⁻ as well as their formation temperatures. Assuming equilibrium, the temperature of dolomite-magnetite precipitation in the Ryugu A0058 sample analysed previously is estimated to be 37 ± 10 °C (ref. 8), while the inferred alteration temperatures of CI chondrites range up to 150 °C (ref. 22). The equilibrium O isotopic fractionation between water and calcite leads to the enhancement of δ¹⁸O values in calcite relative to water by approximately +38‰, +28‰ and +13‰ at 0, 40 and 150 °C, respectively²³. Thus, the δ¹⁸O variation of -22‰ observed in the Ryugu calcite (Fig. 2) is potentially explained by formation temperatures that varied from 0 to 150 °C assuming a fixed δ¹⁸O value of water.

However, this argument does not necessarily mean that variable formation temperature is the sole explanation for the observed δ¹⁸O variation of the calcite. Indeed, the lack of a simple correlation between δ¹⁸O and δ¹³C values (Fig. 3) implies that variable formation temperatures alone cannot explain the observed δ¹⁸O and δ¹³C variations because the δ¹⁸O and δ¹³C values of carbonates should co-vary with their formation temperatures¹⁸. Rather, it seems likely that the δ¹⁸O value of water and/or the δ¹³C value of CO₃²⁻ varied spatially and/or temporally. A previous study used clumped isotope thermometry for carbonates in Mighei-type (CM) chondrites and demonstrated that the δ¹⁸O values of water are variable between samples¹⁶.
<table><caption>Table 1 | δ¹⁸O, δ¹⁷O, Δ¹⁷O and δ¹³C values of carbonates in Ryugu and Ivuna samples</caption><thead><tr><th>Samples</th><th>δ¹⁸O (‰)</th><th>2σ</th><th>δ¹⁷O (‰)</th><th>2σ</th><th>Δ¹⁷O (‰)</th><th>2σ</th><th>δ¹³C (‰)</th><th>2σ</th></tr></thead><tbody><tr><td colspan="9"><strong>Ryugu C0002, calcite</strong></td></tr><tr><td>Calcite#1a<sup>a</sup></td><td>23.5</td><td>0.8</td><td>12.5</td><td>1.6</td><td>0.2</td><td>1.5</td><td>98.1</td><td>4.9</td></tr><tr><td>Calcite#1b<sup>a</sup></td><td>23.5</td><td>0.8</td><td>12.5</td><td>1.6</td><td>0.2</td><td>1.5</td><td>92.7</td><td>4.8</td></tr><tr><td>Calcite#1 average</td><td>23.5</td><td>0.8</td><td>12.5</td><td>1.6</td><td>0.2</td><td>1.5</td><td>95.4</td><td>3.4</td></tr><tr><td>Calcite#2</td><td>38.4</td><td>0.6</td><td>22.2</td><td>1.1</td><td>2.2</td><td>1.1</td><td>101.7</td><td>4.8</td></tr><tr><td>Calcite#3-1</td><td>44.3</td><td>0.8</td><td>24.6</td><td>1.4</td><td>1.6</td><td>1.5</td><td>89.9</td><td>5.2</td></tr><tr><td>Calcite#3-2a<sup>a</sup></td><td>46.3</td><td>0.8</td><td>25.9</td><td>1.6</td><td>1.9</td><td>1.7</td><td>76.2</td><td>5.7</td></tr><tr><td>Calcite#3-2b<sup>a</sup></td><td>46.3</td><td>0.8</td><td>25.9</td><td>1.6</td><td>1.9</td><td>1.7</td><td>84.5</td><td>6.5</td></tr><tr><td>Calcite#3-2c<sup>a</sup></td><td>46.3</td><td>0.8</td><td>25.9</td><td>1.6</td><td>1.9</td><td>1.7</td><td>81.0</td><td>5.9</td></tr><tr><td>Calcite#3-2 average</td><td>46.3</td><td>0.8</td><td>25.9</td><td>1.6</td><td>1.9</td><td>1.7</td><td>80.6</td><td>3.5</td></tr><tr><td>Calcite#3-3</td><td>44.5</td><td>1.2</td><td>24.9</td><td>1.5</td><td>1.8</td><td>1.4</td><td>84.1</td><td>6.2</td></tr><tr><td>Calcite#3-4</td><td>42.9</td><td>1.2</td><td>23.3</td><td>1.3</td><td>1.0</td><td>1.4</td><td>87.0</td><td>5.6</td></tr><tr><td>Calcite#4-1</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td>n.d.</td><td></td><td>88.2</td><td>5.2</td></tr><tr><td>Calcite#4-2</td><td>n.d.</td><td></td><td>n.d.</td><td></td><td>n.d.</td><td></td><td>86.1</td><td>4.8</td></tr><tr><td>Calcite#5a<sup>a</sup></td><td>41.1</td><td>1.2</td><td>24.2</td><td>1.4</td><td>2.9</td><td>1.4</td><td>93.9</td><td>4.8</td></tr><tr><td>Calcite#5b<sup>a</sup></td><td>40.9</td><td>1.2</td><td>23.2</td><td>1.3</td><td>1.9</td><td>1.4</td><td>93.9</td><td>4.8</td></tr><tr><td>Calcite#5 average</td><td>41.0</td><td>0.8</td><td>23.7</td><td>1.0</td><td>2.4</td><td>1.0</td><td>93.9</td><td>4.8</td></tr><tr><td colspan="9"><strong>Ivuna, calcite</strong></td></tr><tr><td>Calcite#1</td><td>41.6</td><td>1.0</td><td>22.5</td><td>1.3</td><td>0.9</td><td>1.4</td><td>101.5</td><td>4.1</td></tr><tr><td>Calcite#2</td><td>41.5</td><td>1.0</td><td>23.4</td><td>1.3</td><td>1.8</td><td>1.4</td><td>104.8</td><td>4.8</td></tr><tr><td>Calcite#3</td><td>40.7</td><td>1.1</td><td>21.9</td><td>1.6</td><td>0.8</td><td>1.4</td><td>107.1</td><td>4.7</td></tr><tr><td>Calcite#5</td><td>33.9</td><td>1.0</td><td>20.0</td><td>1.3</td><td>2.4</td><td>1.4</td><td>81.4</td><td>3.8</td></tr><tr><td>Calcite#6</td><td>41.1</td><td>1.0</td><td>24.0</td><td>1.4</td><td>2.6</td><td>1.4</td><td>103.9</td><td>6.0</td></tr><tr><td>Calcite#7</td><td>38.0</td><td>1.0</td><td>20.2</td><td>1.4</td><td>0.4</td><td>1.4</td><td>98.0</td><td>5.1</td></tr><tr><td>Calcite#8-1</td><td>37.6</td><td>1.0</td><td>20.6</td><td>1.3</td><td>1.1</td><td>1.4</td><td>79.5</td><td>4.4</td></tr><tr><td>Calcite#8-2</td><td>37.0</td><td>1.0</td><td>20.0</td><td>1.3</td><td>0.8</td><td>1.4</td><td>89.0</td><td>5.2</td></tr><tr><td>Calcite#9a<sup>a</sup></td><td>36.0</td><td>1.0</td><td>20.2</td><td>1.3</td><td>1.5</td><td>1.4</td><td>106.6</td><td>5.8</td></tr><tr><td>Calcite#9b<sup>a</sup></td><td>36.0</td><td>1.0</td><td>20.2</td><td>1.3</td><td>1.5</td><td>1.4</td><td>109.9</td><td>5.1</td></tr><tr><td>Calcite#9c<sup>a</sup></td><td>36.0</td><td>1.0</td><td>20.2</td><td>1.3</td><td>1.5</td><td>1.4</td><td>106.7</td><td>4.8</td></tr><tr><td>Calcite#9 average</td><td>36.0</td><td>1.0</td><td>20.2</td><td>1.3</td><td>1.5</td><td>1.4</td><td>107.8</td><td>3.0</td></tr><tr><td>Calcite#10a<sup>a</sup></td><td>43.6</td><td>0.6</td><td>22.3</td><td>1.5</td><td>-0.4</td><td>1.4</td><td>65.4</td><td>4.7</td></tr><tr><td>Calcite#10b<sup>a</sup></td><td>40.9</td><td>0.7</td><td>21.2</td><td>1.2</td><td>-0.1</td><td>1.2</td><td>65.4</td><td>4.7</td></tr><tr><td>Calcite#10 average</td><td>42.2</td><td>0.5</td><td>21.7</td><td>1.0</td><td>-0.3</td><td>0.9</td><td>65.4</td><td>4.7</td></tr><tr><td colspan="9"><strong>Ryugu A0058, dolomite</strong></td></tr><tr><td>Dolomite#5<sup>a</sup></td><td>(27.0)</td><td>(0.9)</td><td>(14.2)</td><td>(1.0)</td><td>(0.2)</td><td>(0.9)</td><td>67.7</td><td>1.5</td></tr><tr><td>Dolomite#5<sup>a</sup></td><td>(27.0)</td><td>(0.9)</td><td>(14.2)</td><td>(1.0)</td><td>(0.2)</td><td>(0.9)</td><td>67.2</td><td>1.2</td></tr><tr><td>Dolomite#5 average</td><td>(27.0)</td><td>(0.9)</td><td>(14.2)</td><td>(1.0)</td><td>(0.2)</td><td>(0.9)</td><td>67.5</td><td>1.0</td></tr><tr><td>Dolomite#2</td><td>(29.0)</td><td>(1.0)</td><td>(15.0)</td><td>(1.0)</td><td>(0.0)</td><td>(0.9)</td><td>74.8</td><td>1.5</td></tr><tr><td>Dolomite#4</td><td>(30.0)</td><td>(0.9)</td><td>(15.8)</td><td>(1.0)</td><td>(0.2)</td><td>(0.9)</td><td>72.1</td><td>1.3</td></tr><tr><td>Dolomite#1</td><td>(29.9)</td><td>(0.9)</td><td>(14.8)</td><td>(1.0)</td><td>(-0.7)</td><td>(0.9)</td><td>71.3</td><td>1.2</td></tr><tr><td>Dolomite#3</td><td>(29.5)</td><td>(0.8)</td><td>(15.7)</td><td>(1.0)</td><td>(0.4)</td><td>(0.9)</td><td>74.5</td><td>1.5</td></tr><tr><td colspan="9"><strong>Ryugu C0002, dolomite</strong></td></tr><tr><td>Dolomite#3a<sup>a</sup></td><td>27.3</td><td>1.0</td><td>15.2</td><td>1.3</td><td>1.0</td><td>0.9</td><td>72.1</td><td>1.3</td></tr><tr><td>Dolomite#3b<sup>a</sup></td><td>28.1</td><td>1.0</td><td>14.7</td><td>1.3</td><td>0.1</td><td>0.9</td><td>72.1</td><td>1.3</td></tr><tr><td>Dolomite#3 average</td><td>27.7</td><td>0.7</td><td>14.9</td><td>0.9</td><td>0.6</td><td>0.6</td><td>72.1</td><td>1.3</td></tr><tr><td>Dolomite#4</td><td>28.3</td><td>1.0</td><td>15.3</td><td>1.3</td><td>0.6</td><td>0.9</td><td>70.1</td><td>1.3</td></tr><tr><td>Dolomite#1a<sup>a</sup></td><td>28.6</td><td>1.0</td><td>15.6</td><td>1.3</td><td>0.7</td><td>0.9</td><td>68.1</td><td>1.3</td></tr><tr><td>Dolomite#1b<sup>a</sup></td><td>28.7</td><td>1.0</td><td>15.1</td><td>1.3</td><td>0.2</td><td>0.9</td><td>68.1</td><td>1.3</td></tr></tbody></table>
<table><caption>Table 1 (continued) | δ¹⁸O, δ¹⁷O, Δ¹⁷O and δ¹³C values of carbonates in Ryugu and Ivuna samples</caption><thead><tr><th>Samples</th><th>δ¹⁸O (‰)</th><th>2σ</th><th>δ¹⁷O (‰)</th><th>2σ</th><th>Δ¹⁷O (‰)</th><th>2σ</th><th>δ¹³C (‰)</th><th>2σ</th></tr></thead><tbody><tr><td>Dolomite#1 average</td><td>28.7</td><td>0.7</td><td>15.4</td><td>0.9</td><td>0.5</td><td>0.6</td><td>68.1</td><td>1.3</td></tr><tr><td>Dolomite#2a<sup>a</sup></td><td>26.7</td><td>1.0</td><td>14.1</td><td>1.3</td><td>0.2</td><td>0.9</td><td>69.4</td><td>1.3</td></tr><tr><td>Dolomite#2b<sup>a</sup></td><td>27.7</td><td>1.0</td><td>14.1</td><td>1.3</td><td>-0.3</td><td>0.9</td><td>69.5</td><td>1.3</td></tr><tr><td>Dolomite#2 average</td><td>27.2</td><td>0.7</td><td>14.1</td><td>0.9</td><td>0.0</td><td>0.6</td><td>69.4</td><td>0.9</td></tr><tr><td><strong>Ivuna, dolomite</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Dolomite#1</td><td>(29.0)</td><td>(0.7)</td><td>(15.6)</td><td>(0.8)</td><td>(0.5)</td><td>(0.9)</td><td>68.1</td><td>1.3</td></tr><tr><td>Dolomite#7</td><td>(30.8)</td><td>(0.7)</td><td>(16.0)</td><td>(0.8)</td><td>(0.0)</td><td>(0.7)</td><td>68.6</td><td>1.1</td></tr><tr><td>Dolomite#6</td><td>(29.1)</td><td>(0.8)</td><td>(15.3)</td><td>(0.8)</td><td>(0.2)</td><td>(0.9)</td><td>72.3</td><td>1.2</td></tr><tr><td>Dolomite#5</td><td>(28.8)</td><td>(0.8)</td><td>(16.3)</td><td>(0.8)</td><td>(1.3)</td><td>(0.9)</td><td>67.2</td><td>1.2</td></tr><tr><td>Dolomite#4</td><td>(26.5)</td><td>(0.9)</td><td>(14.4)</td><td>(0.9)</td><td>(0.7)</td><td>(1.0)</td><td>68.3</td><td>1.3</td></tr><tr><td>Dolomite#2a<sup>a</sup></td><td>(29.6)</td><td>(0.9)</td><td>(15.1)</td><td>(0.9)</td><td>(-0.3)</td><td>(0.8)</td><td>67.4</td><td>1.3</td></tr><tr><td>Dolomite#2b<sup>a</sup></td><td>(29.6)</td><td>(0.9)</td><td>(15.1)</td><td>(0.9)</td><td>(-0.3)</td><td>(0.8)</td><td>65.6</td><td>1.3</td></tr><tr><td>Dolomite#2c<sup>a</sup></td><td>(29.6)</td><td>(0.9)</td><td>(15.1)</td><td>(0.9)</td><td>(-0.3)</td><td>(0.8)</td><td>67.4</td><td>1.3</td></tr><tr><td>Dolomite#2 average</td><td>(29.6)</td><td>(0.9)</td><td>(15.1)</td><td>(0.9)</td><td>(-0.3)</td><td>(0.8)</td><td>66.8</td><td>0.7</td></tr><tr><td>Dolomite#3</td><td>(29.2)</td><td>(1.0)</td><td>(15.4)</td><td>(1.0)</td><td>(0.2)</td><td>(1.1)</td><td>68.8</td><td>1.2</td></tr></tbody></table>

The O isotope compositions of the Ryugu A0058 and Ivuna dolomite, shown in parentheses, are taken from Yokoyama et al.⁸. The 2σ errors are either external reproducibility (2s.d.) of standard measurements or internal precision (2s.e.) of the data within single measurements, whichever is larger (Supplementary Table 1). n.d., not determined. ⁹For single grains that were analysed multiple times, we averaged the data and calculated the corresponding uncertainties by propagating individual errors.

Mass-balance calculations²⁴,²⁵ and the O isotope composition of the putative, early Solar System water²⁶,²⁷ suggest that before the onset of alteration, water in the CI and other carbonaceous chondrites had a much higher Δ¹⁷O value, and possibly δ¹⁸O value, than the anhydrous silicates²⁸. Thus, as alteration progressed, the Δ¹⁷O value, and possibly the δ¹⁸O value, of the altering fluid would have decreased²⁴,²⁵. The Δ¹⁷O values of the carbonates and fluid will be identical at equilibrium, and thus, the Δ¹⁷O values of the carbonates are a measure of the degree of progress of water-rock interactions. The Δ¹⁷O values of the dolomite in the Ryugu and Ivuna samples are systematically lower than those of the calcite, and the Ryugu and Ivuna carbonates show resolvable Δ¹⁷O variations beyond uncertainties (Fig. 2b). Therefore, the calcite with systematically higher Δ¹⁷O values formed from less ‘evolved’ fluids and crystallized earlier than the dolomite. The fact that calcite is more prevalent in less-altered areas, as shown by primary anhydrous silicates, suggests that Ca was more easily leached than Mg during incipient aqueous alteration, allowing the formation of calcite before dolomite.

Like O isotopes, the C isotope compositions of carbonates are determined not only by their formation temperatures but also by the δ¹³C values of dissolved CO₃²⁻. The C isotope composition of CO₃²⁻ could have varied due to (1) Rayleigh-type isotopic fractionation as a result of the preferential escape of ¹²C-rich gaseous species such as CH₄ (ref. 16), (2) the progressive formation of carbonates, that is, fractional crystallization, (3) the mixing of two or more C reservoirs with distinct δ¹³C values that supplied CO₃²⁻ (ref. 19) or (4) the change in the chemical speciation of the C-bearing gaseous species, such as CO₂, CO and CH₄, due to varying O and H partial pressures. Rayleigh-type isotopic fractionation was not the primary mechanism for producing the observed δ¹³C variation because it would have resulted in higher δ¹³C values in the dolomite, which formed from the more-evolved fluids than the calcite, whereas the opposite is observed. The influence of fractional crystallization during carbonate formation was also minimal because the rare calcite that formed early presumably when a larger CO₃²⁻ pool was available displays a larger δ¹³C variation than the more common dolomite, which is opposite to what one would expect in this scenario.

A previous study advocated that the observed δ¹³C variation resulted from the mixing of C reservoirs with distinct δ¹³C values reflecting the spatially heterogeneous distribution of different C reservoirs¹⁹. Possible C reservoirs include C-bearing gaseous species such as CO₂,

CO and CH₄, originally accreted in ices, and organic matter, and the previous study invoked CO₂-bearing ices as a ¹³C-rich reservoir¹⁹. However, the mixing timescales of gaseous species in the Ryugu/CI parent asteroids would have been short unless the permeability was extremely low. Therefore, the spatially heterogeneous distribution of C reservoirs would have not persisted for long. Furthermore, it is unclear why the calcite has commonly higher δ¹³C values than the dolomite, and calcite with lower δ¹³C values is generally absent.

## Temporal change in oxygen fugacity and gaseous species

Instead of spatial heterogeneity, the temporal variation in the δ¹³C values of the C reservoirs and their chemical speciation probably occurred due to a change in O partial pressure or, more precisely, O fugacity (fO₂), which is O partial pressure corrected for nonideal gas behaviour. Oxygen fugacity varied along with the production of H₂ via the oxidation of Fe in metal and silicates by H₂O and the subsequent escape of H₂ from the system, perhaps by diffusion or by making fractures in the parent body²⁹. Thus, fO₂ was determined by the relative rates of the production and escape of H₂, in this scenario, fO₂ would have at first decreased and then increased. In the case of CM chondrites, the amount of Fe³⁺ in their matrices increases with increasing alteration³⁰, which is in line with this scenario.

The Fe and Mn abundances of terrestrial carbonates that reflect the Fe²⁺ and Mn²⁺ activities in fluids have been used to infer the redox conditions under which they precipitated³¹. However, the Fe and Mn abundances of Ryugu and Ivuna carbonates were probably controlled not only by redox conditions but also by the amounts of these cations leached from primary minerals during the progressive aqueous alteration³². Therefore, the zoning of Fe and Mn observed in the dolomite grains may not be a direct proxy for the temporal change in fO₂. Instead, here we propose that the δ¹³C values of Ryugu and Ivuna carbonates represent a record of such temporal fO₂ variation.

To see how the δ¹³C values of carbonates will change with varying fO₂, we consider a rather simple model, where gaseous CO₂ and CO and carbonates (and dissolved CO₂, HCO₃⁻ and CO₃²⁻) are in C isotopic equilibrium, and the CO₂/CO ratio increases, corresponding to an increase in fO₂. In cometary ices, CO₂ and CO are the most abundant C-bearing chemical species³³, and the Ryugu/CI parent bodies would
<figure><img src="image_6.png" /></figure>

<figure><img src="image_7.png" /><figcaption>**Fig. 2 |** Oxygen isotope compositions of the calcite and dolomite in Ryugu and Ivuna samples. The whole-rock O isotope compositions of the Ryugu A0107 and C0002 samples⁸ are also shown for reference. The changes in O isotope compositions due to variable formation temperatures and water-rock interaction are illustrated by arrows. **a**, Oxygen three-isotope plot showing δ¹⁸O versus δ¹⁷O values. **b**, δ¹⁸O versus Δ¹⁷O values. Data are presented as mean values ± 2σ errors, which are either external reproducibility (2 s.d., *N* = 6–20 depending on the measurement sessions) of standard measurements or internal precision (2 s.e.) of the data within single measurements, whichever is larger. TFL, terrestrial fractionation line, defined as δ¹⁷O = 0.52 × δ¹⁸O.</figcaption></figure>

have accreted substantial amounts of CO₂ and CO-bearing ices if they
formed in the distal Solar System<sup>11,28</sup>. At the earliest stage of aqueous
alteration, the CO₂/CO ratio may have been characterized by that of
the accreted ices, which may be around unity or higher as observed
for cometary ices<sup>33,34</sup>. We assume that the δ¹³C value of the bulk gas
(CO₂ + CO), δ¹³C<sub>bulk</sub>, is constant regardless of the CO₂/CO ratio. Then
the δ¹³C<sub>bulk</sub> is given by mass balance as δ¹³C<sub>bulk</sub> = xδ¹³C<sub>CO₂</sub> + (1-x) δ¹³C<sub>CO</sub>,
where x is the mole fraction of CO₂ defined by x = CO₂/(CO₂ + CO) and
δ¹³C<sub>CO₂</sub> and δ¹³C<sub>CO</sub> are the δ¹³C values of CO₂ and CO, respectively.
Thus, using Δ defined by δ¹³C<sub>CO₂</sub> - δ¹³C<sub>CO</sub>, δ¹³C<sub>CO₂</sub> and δ¹³C<sub>CO</sub> are given
by δ¹³C<sub>CO₂</sub> = δ¹³C<sub>bulk</sub> + (1-x) Δ and δ¹³C<sub>CO</sub> = δ¹³C<sub>bulk</sub> - x Δ, respectively.
The Δ value is approximated by 1,000 lnα, where α is the C isotopic frac-
tionation factor between CO₂ and CO defined by (¹³C/¹²C)CO₂/(¹³C/¹²C)CO
and is positive at all temperatures (+93‰, +76‰ and +48‰ at 0, 40 and

<figure><img src="image_8.png" /><figcaption>**Fig. 3 |** Comparison between C and O isotope compositions of the calcite and dolomite in Ryugu and Ivuna samples. The values of carbonates measured for whole-rock CI chondrites (Ivuna and Orgueil)¹⁸ are also shown for reference. The changes in C and O isotope compositions due to variable formation temperatures and O fugacity, and water-rock interaction, are illustrated by arrows. The calcite shows much larger variations in both C and O isotope compositions than the dolomite. Note that no simple correlation between δ¹³C and δ¹⁸O values can be seen. Data are presented as mean values ± 2σ errors, which are either external reproducibility (2 s.d., *N* = 6–20 depending on the measurement sessions) of standard measurements or internal precision (2 s.e.) of the data within single measurements, whichever is larger.</figcaption></figure>

150 °C, respectively<sup>21,35</sup>). Therefore, with increasing x (and fO<sub>2</sub>), the δ<sup>13</sup>C values of both CO<sub>2</sub> and CO will decrease monotonically, and the δ<sup>13</sup>C values of carbonates will also decrease. Thus, in this simple model, a δ<sup>13</sup>C variation of carbonates comparable to the observation (~40‰) is expected if the x value varies from 0.5 to 1 (Fig. 3). The redox states evolved within a few million years after the birth of the Solar System as inferred from the <sup>53</sup>Mn–<sup>53</sup>Cr chronometry of Ryugu/CI dolomite<sup>8,10,20,36</sup>. The presence of CH<sub>4</sub>, another possible reducing gas, does not change this conclusion because the behaviour of CH<sub>4</sub> in terms of C isotopic fractionation against CO<sub>2</sub> is quite similar to that of CO (ref. 21). Our model where CO<sub>2</sub> and CO are in C isotopic equilibrium at low temperatures requires the presence of as yet unidentified processes/catalysts. The isotopic equilibrium in thermally matured natural gases has been recently discussed<sup>37</sup>.

The characteristic δ¹³C values of Ryugu and Ivuna carbonates are not observed for other aqueously altered meteorites such as CM chondrites and the ungrouped carbonaceous chondrite Tagish Lake<sup>19,38,39</sup>, suggesting that the redox conditions as well as the thermal history and accreted materials are unique to their parent bodies. For example, the δ¹³C values of CM calcite are variable like Ryugu/CI calcite, but the highest reported value (+80‰) in CM calcite is lower than that of Ryugu/CI calcite (+108‰) (Extended Data Fig. 2). Furthermore, the δ¹³C values of CM dolomite are also variable from ~40‰ to ~60‰, which is in contrast to the homogeneous δ¹³C values observed for Ryugu/CI dolomite. Iron in CI chondrites is dominated by octahedral Fe³⁺ and indicates more-oxidized conditions than CM chondrites³⁰, which is consistent with the homogeneous δ¹³C values of Ryugu/CI dolomite.
Because the δ¹³C values of the Ryugu/CI dolomite are commonly lower and more homogeneous than those of the calcite, the dolomite probably formed at higher fO₂ and/or temperature than the calcite. This, combined with the O isotope signatures of the carbonates, implies that the calcite formed during prograde alteration over wide ranges of fO₂ and temperature, whereas the dolomite formed later during retrograde cooling when the aqueous fluids and silicates approached O isotopic equilibrium. Therefore, when the dolomite formed, the fO₂ was probably high enough that the major gaseous C reservoir was CO₂ (that is, x ~ 1). If correct, the δ¹³C value of CO₂ in C isotopic equilibrium with the dolomite, which has the average δ¹³C value of ~+70‰, would equal the δ¹³C value of the bulk gas. Using the C isotopic fractionation factor between dolomite and CO₂, (¹³C/¹²C)₉₅%ₗ/(¹³C/¹²C)₉₅%ₘ, of ~1.0092 at 40 °C or ~0.9951 at 150 °C (refs. 40,41), the δ¹³Cₙₘₗ = (δ¹³C₉₅%ₗ - δ¹³C₉₅%ₘ) value is estimated to be between +61‰ and +75‰.

The preceding simple model implies a δ¹³C value higher than +60‰ for C-bearing molecules originally accreted in ices. Such ¹³C-rich compositions are not common among Solar System materials other than meteoritic carbonates and water-soluble organic compounds⁴². A similar level of ¹³C-enrichment (δ¹³C = +65 ± 51‰), albeit with large uncertainty, has been reported for CO₂ in the coma of 67P/Churyumov-Gerasimenko⁴³. The mechanism to produce ¹³C-rich compositions is not well understood, but self-shielding during CO photodissociation in the solar nebula or the parent molecular cloud of the Solar System is a possible mechanism¹⁴⁻¹⁶. Thus, the inferred ¹³C-rich composition of C-bearing molecular ices would have resulted from such physicochemical reactions⁴⁷,⁴⁸, and we concluded that the Ryugu/CI parent bodies accreted materials that originated from these cold environments.

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Wataru Fujiya Ⓓ1, Noriyuki Kawasaki Ⓓ2, Kazuhide Nagashima Ⓓ3, Naoya Sakamoto Ⓓ4, Conel M. O'D. Alexander⁵, Noriko T. Kita⁶, Kouki Kitajima Ⓓ6, Yoshinari Abe Ⓓ7, Jérôme Aléon Ⓓ8, Sachiko Amari Ⓓ9,¹⁰, Yuri Amelin¹¹, Ken-ichi Bajo Ⓓ2, Martin Bizzarro¹², Audrey Bouvier Ⓓ13, Richard W. Carlson Ⓓ5, Marc Chaussidon¹⁴, Byeon-Gak Choi Ⓓ15, Nicolas Dauphas Ⓓ16, Andrew M. Davis¹⁶, Tommaso Di Rocco Ⓓ17, Ryota Fukai Ⓓ18, Ikshu Gautam Ⓓ19, Makiko K. Haba Ⓓ19, Yuki Hibiya Ⓓ20, Hiroshi Hidaka Ⓓ21, Hisashi Homma Ⓓ22, Peter Hoppe Ⓓ23, Gary R. Huss Ⓓ3, Kiyohiro Ichida Ⓓ24, Tsuyoshi Iizuka Ⓓ25, Trevor R. Ireland Ⓓ26, Akira Ishikawa Ⓓ19, Shoichi Itoh Ⓓ27, Thorsten Kleine Ⓓ28, Shintaro Komatani²⁴, Alexander N. Krot Ⓓ3, Ming-Chang Liu²⁹,³⁰, Yuki Masuda Ⓓ19, Kevin D. McKeegan²⁹, Mayu Morita Ⓓ24, Kazuko Motomura³¹, Frédéric Moynier Ⓓ14, Izumi Nakai³², Ann Nguyen Ⓓ33, Larry Nittler Ⓓ5, Morihiko Onose²⁴, Andreas Pack¹⁷, Changkun Park Ⓓ34, Laurette Piani³⁵, Liping Qin Ⓓ36, Sara S. Russell³⁷, Maria Schönbächler Ⓓ38, Lauren Tafla²⁹, Haolan Tang²⁹, Kentaro Terada Ⓓ39, Yasuko Terada⁴⁰, Tomohiro Usui Ⓓ18, Sohei Wada Ⓓ2, Meenakshi Wadhwa⁴¹, Richard J. Walker Ⓓ42, Katsuyuki Yamashita⁴³, Qing-Zhu Yin Ⓓ44, Tetsuya Yokoyama Ⓓ19, Shigekazu Yoneda Ⓓ45, Edward D. Young Ⓓ29, Hiroharu Yui⁴⁶, Ai-Cheng Zhang Ⓓ47, Tomoki Nakamura⁴⁸, Hiroshi Naraoka⁴⁹, Takaaki Noguchi Ⓓ27, Ryuji Okazaki⁴⁹, Kanako Sakamoto Ⓓ18, Hikaru Yabuta⁵⁰, Masanao Abe¹⁸, Akiko Miyazaki Ⓓ18, Aiko Nakato Ⓓ18, Masahiro Nishimura¹⁸, Tatsuaki Okada¹⁸, Toru Yada Ⓓ18, Kasumi Yogata¹⁸, Satoru Nakazawa Ⓓ18, Takanao Saiki¹⁸, Satoshi Tanaka¹⁸, Fuyuto Terui⁵¹, Yuichi Tsuda¹⁸, Sei-ichiro Watanabe Ⓓ21, Makoto Yoshikawa¹⁸, Shogo Tachibana Ⓓ52 & Hisayoshi Yurimoto Ⓓ2

<footnote n="1">¹Faculty of Science, Ibaraki University, Mito, Japan. ²Department of Natural History Sciences, Hokkaido University, Sapporo, Japan. ³Hawai'i Institute of Geophysics and Planetology, University of Hawai'i at Mānoa, Honolulu, HI, USA. ⁴Isotope Imaging Laboratory, Creative Research Institution, Hokkaido University, Sapporo, Japan. ⁵Earth and Planets Laboratory, Carnegie Institution for Science, Washington, DC, USA. ⁶Geoscience, University of Wisconsin-Madison, Madison, WI, USA. ⁷Graduate School of Engineering Materials Science and Engineering, Tokyo Denki University, Tokyo, Japan. ⁸Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie, Sorbonne Université, Museum National d'Histoire Naturelle, CNRS UMR 7590, IRD, Paris, France. ⁹McDonnell Center for the Space Sciences and Physics Department, Washington University, St Louis, MO, USA. ¹⁰Geochemical Research Center,</footnote>
The University of Tokyo, Tokyo, Japan. ¹¹Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou, China. ¹²Centre for Star and Planet Formation, GLOBE Institute, University of Copenhagen, Copenhagen, Denmark. ¹³Bayerisches Geoinstitut, Universität Bayreuth, Bayreuth, Germany. ¹⁴Université Paris Cité, Institut de Physique du Globe de Paris, CNRS, Paris, France. ¹⁵Department of Earth Science Education, Seoul National University, Seoul, Republic of Korea. ¹⁶Department of the Geophysical Sciences and Enrico Fermi Institute, The University of Chicago, Chicago, IL, USA. ¹⁷Faculty of Geosciences and Geography, University of Göttingen, Göttingen, Germany. ¹⁸ISAS/JSEC, JAXA, Sagamihara, Japan. ¹⁹Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Tokyo, Japan. ²⁰Research Center for Advanced Science and Technology, The University of Tokyo, Tokyo, Japan. ²¹Earth and Planetary Sciences, Nagoya University, Nagoya, Japan. ²²Osaka Application Laboratory, SBUWDX, Rigaku Corporation, Osaka, Japan. ²³Max Planck Institute for Chemistry, Mainz, Germany. ²⁴Analytical Technology, Horiba Techno Service Co., Ltd, Kyoto, Japan. ²⁵Earth and Planetary Science, The University of Tokyo, Tokyo, Japan. ²⁶School of Earth and Environmental Sciences, The University of Queensland, St Lucia, Queensland, Australia. ²⁷Earth and Planetary Sciences, Kyoto University, Kyoto, Japan. ²⁸Max Planck Institute for Solar System Research, Göttingen, Germany. ²⁹Earth, Planetary, and Space Sciences, UCLA, Los Angeles, CA, USA. ³⁰Lawrence Livermore National Laboratory, Livermore, CA, USA. ³¹Thermal Analysis, Rigaku Corporation, Tokyo, Japan. ³²Applied Chemistry, Tokyo University of Science, Tokyo, Japan. ³³Astromaterials Research and Exploration Science, NASA Johnson Space Center, Houston, TX, USA. ³⁴Earth-System Sciences, Korea Polar Research Institute, Incheon, Korea. ³⁵Centre de Recherches Pétrographiques et Géochimiques, CNRS - Université de Lorraine, Nancy, France. ³⁶School of Earth and Space Sciences, University of Science and Technology of China, Hefei, China. ³⁷Department of Earth Sciences, Natural History Museum, London, UK. ³⁸Institute for Geochemistry and Petrology, Department of Earth Sciences, ETH Zurich, Zurich, Switzerland. ³⁹Earth and Space Science, Osaka University, Osaka, Japan. ⁴⁰Spectroscopy and Imaging, Japan Synchrotron Radiation Research Institute, Hyogo, Japan. ⁴¹School of Earth and Space Exploration, Arizona State University, Tempe, AZ, USA. ⁴²Geology, University of Maryland, College Park, MD, USA. ⁴³Graduate School of Natural Science and Technology, Okayama University, Okayama, Japan. ⁴⁴Earth and Planetary Sciences, University of California, Davis, CA, USA. ⁴⁵Science and Engineering, National Museum of Nature and Science, Tsukuba, Japan. ⁴⁶Chemistry, Tokyo University of Science, Tokyo, Japan. ⁴⁷School of Earth Sciences and Engineering, Nanjing University, Nanjing, China.

⁴⁸Department of Earth Science, Tohoku University, Sendai, Japan. ⁴⁹Department of Earth and Planetary Sciences, Kyushu University, Fukuoka, Japan.

⁵⁰Earth and Planetary Systems Science Program, Hiroshima University, Higashi-Hiroshima, Japan. ⁵¹Kanagawa Institute of Technology, Atsugi, Japan.

⁵²UTokyo Organization for Planetary and Space Science, University of Tokyo, Tokyo, Japan. e-mail: wataru.fujiya.sci@vc.ibaraki.ac.jp
# Methods

## Isotope measurement using ion microprobe

We produced the polished sections of Ryugu samples A0058-C1001 and C0002-C1001 and Ivuna embedded in epoxy<sup>8,28</sup>. The polished sections were coated with a thin (~5 nm) gold film using a Leica EM ACE600 coater at Hokkaido University for backscattered electron and X-ray imaging as well as elemental analysis, which were conducted before in-situ O and C isotope measurements. We observed their mineralogy and petrology and located carbonate grains using a field-emission scanning electron microscope (JEOL JSM-7000F) equipped with an energy-dispersive X-ray spectrometer (Oxford X-Max 150) at Hokkaido University. The beam currents were ~2 nA and ~1 nA for the X-ray mapping and quantitative analysis, respectively. Quantitative calculations were conducted using Oxford AZtec software.

We selected five and four dolomite grains from the Ryugu samples A0058 and C0002, respectively, and nine calcite grains from C0002 for isotope analyses. We also analysed seven dolomite grains and ten calcite grains from Ivuna. Before the isotope analyses, the samples were coated again with an additional thin (~65 nm) gold film.

We conducted the O and C isotope measurements of the selected carbonate grains using secondary ion mass spectrometry (SIMS; CAMECAims-1280HR) at Hokkaido University. Instrumental mass fractionation was corrected using the UWC3 calcite standard and a series of dolomite–ankerite standards from the WiscSIMS laboratory for calcite and dolomite, respectively<sup>49–51</sup>. Measurement spots were observed using the field-emission scanning electron microscope after the SIMS measurements, and data from spots with inclusions or overlapping matrix minerals were rejected. The reported uncertainties (2σ) of isotope compositions are the larger of the external reproducibility (2 s.d. of standard measurements) or internal precision (2 s.e. of data within single measurements of unknown samples) (Supplementary Table 1).

The procedures of O isotope measurement of dolomite were described by a previous study<sup>8</sup>. For O isotope measurement of calcite, as previously described for the O isotope measurement of olivine<sup>28</sup>, secondary <sup>16</sup>O<sup>-</sup>, <sup>17</sup>O<sup>-</sup> and <sup>18</sup>O<sup>-</sup> ions produced by a Cs<sup>+</sup> primary ion beam (~30 pA, ~3 µm) were simultaneously collected using Faraday cup (10<sup>11</sup> Ω), electron multiplier (EM) and EM detectors, respectively. Mass-resolving power was >6,000, sufficient to resolve <sup>17</sup>O<sup>-</sup> from <sup>16</sup>OH<sup>-</sup>. The secondary ion intensities of <sup>16</sup>O<sup>-</sup> were 2–3 × 10<sup>7</sup> cps. The measurement time was 240 s. The <sup>16</sup>OH<sup>-</sup> count rate was measured immediately after each measurement, and we made a small tail correction on <sup>17</sup>O<sup>-</sup>; its contribution to <sup>17</sup>O<sup>-</sup> was typically less than ~0.1‰ and up to ~0.5‰ for a few analyses. The typical uncertainties of δ<sup>17</sup>O, δ<sup>18</sup>O and Δ<sup>17</sup>O values were 1.4‰, 0.9‰ and 1.4‰, respectively.

For C isotope measurement of dolomite, secondary <sup>12</sup>C<sup>-</sup> and <sup>13</sup>C<sup>-</sup> ions produced by a Cs<sup>+</sup> primary ion (~50 pA, ~2 µm) were simultaneously collected using Faraday cup (10<sup>12</sup> Ω) and EM detectors, respectively. We scanned the primary ion beam across 1 × 1-µm-sized areas to make the SIMS pits shallower and suppress ratio drifts during measurements. Mass-resolving power was ~4,500, sufficient to resolve <sup>13</sup>C<sup>-</sup> from <sup>12</sup>CH<sup>-</sup>. The secondary ion intensities of <sup>12</sup>C<sup>-</sup> were 5–6 × 10<sup>5</sup> cps. The measurement time was 480 s. The typical uncertainty of δ<sup>13</sup>C values was 1.3‰.

For C isotope measurement of calcite, secondary <sup>12</sup>C<sup>-</sup> and <sup>13</sup>C<sup>-</sup> ions produced by a Cs<sup>+</sup> primary ion (~3 pA, ~1 µm) were simultaneously collected using two EM detectors. We scanned the primary ion beam on 1 × 1-µm-sized areas. Mass-resolving power was ~4,500. The secondary ion intensities of <sup>12</sup>C<sup>-</sup> were ~2 × 10<sup>4</sup> cps. The measurement time was 800 s. The typical uncertainty of δ<sup>13</sup>C values was 5.1‰. The larger uncertainty of the calcite measurement was due to the smaller primary ion beam intensity than for the dolomite measurement to analyse small calcite grains.

# Data availability

All data generated or analysed during this study are included in this published article (and its supplementary information files) and are available via Zenodo (https://doi.org/10.5281/zenodo.7957625). As the

initial analysis of Ryugu samples collected by the Hayabusa2 spacecraft, the specimens analysed in this study were allocated to us by JAXA. The Ivuna specimen used in this study was kindly proved by the Natural History Museum, UK.

# References

49. Kozdon, R., Ushikubo, T., Kita, N. T., Spicuzza, M. & Valley, J. W. Intratest oxygen isotope variability in the planktonic foraminifer *N. pachyderma*: real vs. apparent vital effects by ion microprobe. *Chem. Geol.* **258**, 327–337 (2009).

50. Śliwiński, M. G. et al. Secondary ion mass spectrometry bias on isotope ratios in dolomite–ankerite, part I: δ<sup>18</sup>O matrix effects. *Geostand. Geoanal. Res.* **40**, 157–172 (2015).

51. Śliwiński, M. G. et al. Secondary ion mass spectrometry bias on isotope ratios in dolomite-ankerite, part II: δ<sup>13</sup>C matrix effects. *Geostand. Geoanal. Res.* **40**, 173–184 (2015).

# Acknowledgements

We thank D. Rogers, M. Spicuzza and J. Valley for the preparation of carbonate standard materials for SIMS measurements and A. Tsuchiyama for discussion. Hayabusa2 was developed and built under the leadership of Japan Aerospace Exploration Agency (JAXA), with contributions from the German Aerospace Center (DLR) and the Centre National d'Études Spatiales (CNES), and in collaboration with NASA, and other universities, institutes and companies in Japan. The curation system was developed by JAXA in collaboration with companies in Japan. This research was supported in part by the JSPS KAKENHI grant numbers 19H00725 (W.F.), 20K20934 (W.F. and T.N.) and 22K18722 (N.K.).

# Author contributions

W.F., N.K., K.N., N.S. and H. Yurimoto designed this study and T. Nakamura, H.N., T. Noguchi, R.O., K.S., H. Yabuta, M.A., A.M., A. Nakato, M.N., T.O., T. Yada, K. Yogata, S.N., T.S., S. Tanaka, F.T., Y. Tsuda, S. Watanabe, M.Y. and S. Tachibana supported them. W.F., N.K., K.N., N.S. and H. Yurimoto analysed the samples, and N.T.K. and K.K. assisted the analysis. W.F., N.K., K.N. and H. Yurimoto were involved in data reduction. C.M.O'D.A., Y. Abe, J.A., S.A., Y. Amelin, K.B., M.B., A.B., R.W.C., M.C., B.-G.C., N.D., A.M.D., T.D.R., R.F., I.G., M.K.H., Y.H., H. Hidaka., H. Homma, P.H., G.R.H., K.I., T.I., T.R.I., A.I., S.I., T.K., S.K., A.N.K., M.-C.L., Y.M., K.D.M., M.M., K.M., F.M., I.N., A. Nguyen, L.N., M.O., A.P., C.P., L.P., L.Q., S.S.R., M.S., L.T., H.T., K.T., Y. Terada, T.U., S. Wada, M.W., R.J.W., K. Yamashita, Q.-Z.Y., T. Yokoyama, S.Y., E.D.Y., H. Yui and A.-C.Z. contributed to data interpretation. W.F. wrote the paper with support and approval of all co-authors.

# Competing interests

The authors declare no competing interests.

# Additional information

Extended data is available for this paper at
https://doi.org/10.1038/s41561-023-01226-y.

**Supplementary information** The online version contains supplementary material available at https://doi.org/10.1038/s41561-023-01226-y.

**Correspondence and requests for materials** should be addressed to Wataru Fujiya.

**Peer review information** *Nature Geoscience* thanks John Eiler, Michael Zolensky and Christopher Herd for their contribution to the peer review of this work. Primary Handling Editors: Stefan Lachowycz and Alison Hunt, in collaboration with the *Nature Geoscience* team.

**Reprints and permissions information** is available at
www.nature.com/reprints.
<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>

**Extended Data Fig. 1** | Compositional variation of Ryugu and Ivuna carbonates. (a) Ternary diagram of dolomite and calcite in Ryugu and Ivuna samples (see also Supplementary Table 1). (b-d) Backscattered electron image: BEI (b), Mn Kα₁ X-ray map (c), Fe Kα₁ X-ray map (d) of a dolomite grain in the Ryugu C0002 sample.
<figure><img src="image_13.png" /></figure>

**Extended Data Fig. 2** | Comparison between C and O isotope compositions of carbonates in Ryugu, CI, and CM chondrites. CM chondrite data are taken from Telus et al. (ref. 38). The δ¹³C values of CM calcite are variable like Ryugu/CI calcite, but the highest reported value in CM calcite is lower than that of Ryugu/CI calcite. The δ¹³C and δ¹⁸O values of CM dolomite are also variable, whereas

those of Ryugu/CI dolomite are more homogeneous. Data generated during this study are presented as mean values ± 2σ errors which are either external reproducibility (2 SD, N = 6–20 depending on the measurement sessions) of standard measurements or internal precision (2SE) of the data within single measurements, whichever is larger.