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date: 18 October 2019

Meteorites

Summary and Keywords

For thousands of years, people living in Egypt, China, Greece, Rome, and other parts of the world have been fascinated by shooting stars, which are the light and sound phenomena commonly associated with meteorite impacts. The earliest written record of a meteorite fall is logged by Chinese chroniclers in 687 bce. However, centuries before that, Egyptians had been using “heavenly iron” to make their first iron tools, including a dagger found in King Tutankhamun’s tomb that dates back to the 14th century bce. Even though human beings have a long history of observing meteors and utilizing meteorites, we did not start to recognize their true celestial origin until the Age of Enlightenment. In 1794 German physicist and musician Ernst Chladni was the first to summarize the scientific evidence and to demonstrate that these unique objects are indeed from outside of the Earth. After more than two centuries of joint efforts by countless keen amateur, academic, institutional, and commercial collectors, more than 60,000 meteorites have been catalogued and classified in the Meteoritical Bulletin Database. This number is continually growing, and meteorites are found all over the world, especially in dry and sparsely populated regions such as Antarctica and the Sahara Desert. Although there are thousands of individual meteorites, they can be handily classified into three broad groups by simple examinations of the specimens. The most common type is stony meteorite, which is made of mostly silicate rocks. Iron meteorites are the easiest to be preserved for thousands (or even millions) of years on the Earth’s surface environments, and they are composed of iron and nickel metals. The stony-irons contain roughly the same amount of metals and silicates, and these spectacular meteorites are the favorites of many collectors and museums.

After 200 years, meteoritics (the science of meteorites) has grown out of its infancy and become a vibrant area of research today. The general directions of meteoritic studies are: (1) mineralogy, identifying new minerals or mineral phases that rarely or seldom found on the Earth; (2) petrology, studying the igneous and aqueous textures that give meteorites unique appearances, and providing information about geologic processes on the bodies upon which the meteorites originates; (3) geochemistry, characterizing their major, trace elemental, and isotopic compositions, and conducting interplanetary comparisons; and (4) chronology, dating the ages of the initial crystallization and later on impacting disturbances. Meteorites are the only extraterrestrial samples other than Apollo lunar rocks and Hayabusa asteroid samples that we can directly analyze in laboratories. Through the studies of meteorites, we have quested a vast amount of knowledge about the origin of the Solar System, the nature of the molecular cloud, the solar nebula, the nascent Sun and its planetary bodies including the Earth and its Moon, Mars, and many asteroids. In fact, the 4.6-billion-year age of the whole Solar System is solely defined by the oldest age dated in meteorites, which marked the beginning of everything we appreciate today.

Keywords: meteorite, history, chondrites, achondrites, iron meteorites, classification, isotopes, age of the solar system, presolar grains, building blocks of Earth

Introduction

Meteorites provide a glimpse beyond Earth’s bounds. They are fascinating objects that capture people’s imaginations; they are precious samples used in scientific research. Meteorites and related events have been venerated throughout human history. However, it was only in the 18th century that meteoritics, the study of meteorites, became a scientific discipline. The nascent form of meteoritics was born during the Age of Enlightenment; the formative years for many disciplines of modern science. Meteoritics remained as a minor subject within mineralogy for the first 150 years. Meteorites were collected more as novelties, considered rare types of “minerals,” rather than as subjects for serious scientific research. This attitude changed dramatically in the latter half of 20th century with the advent of nationally funded space exploration programs such as the Apollo missions. The fierce competition between countries during the “Space Age” stimulated the rapid development of the planetary sciences and provided ample funding to build and cultivate the scientific community in relevant fields. As the only extraterrestrial materials (besides the samples that returned directly from extraterrestrial bodies by manned or robotic missions) available to humans that can be directly analyzed, meteorites provide unique opportunities for understanding other planets, the origins of the solar system, and the formation and evolution of planetary bodies. Today, over a thousand trained scientists and enthusiastic amateurs are involved in meteorite studies in its different facets, and hundreds of papers related to meteorites are published in professional journals every year. Meteoritics has established itself as an interdisciplinary science that links mineralogy, geochemistry, planetary sciences, and astrophysics.

The scope of meteoritics is inherently broad. Scientists study meteorites by applying the principles and methods of petrography, mineralogy, chemistry, and physics. In particular, this is done by (1) identifying new minerals or mineral assemblages that formed in unique nebular and planetary environments, which are often considerably different from that on Earth; (2) describing igneous and metamorphic textures that reflect the temperature-pressure-fluid conditions and thermal histories of the parent bodies; (3) determining major and trace elemental compositions of minerals and bulk samples that indicate the petrogenesis and magmatic/hydothermal processes on other planets; (4) analyzing isotopic compositions that reveal meteorite precursors, the genetic relationships among different groups, and the kinetic processes that have caused isotopic fractionation; (5) characterizing the shock structures that have imprinted by impacting events to understand the histories of collisions and the chaotic nature of the early solar system; (6) dating ages of formation and the major events that occurred on other planets or asteroids to establish the timeline of the solar system evolution and history of planetesimals and asteroids; (7) measuring various types of spectra at different wavelengths and linking meteorites to their home asteroids; (8) detecting paleomagnetic features that record the nebular dynamic and physics, and activities of embryo cores; (9) examining meteoritic organic matters to decipher the origin of life; and (10) isolating presolar grains that predated the formation of the Sun to provide information on the sources of building blocks of the solar system.

In this short article, it would be impossible to cover all the topics listed above (let alone those that extend beyond these). The authors will first briefly review the history of meteorite discovery and early scientific investigation. The authors will then summarize the basic knowledge obtained heretofore about meteorites through a discussion of meteorite classifications and the unique characteristics of each major group. Finally, the authors will highlight three additional key research topics: meteorite chronology, presolar grains, and insights into the Earth’s origin. The authors intend to regularly amend the section “Sources of Meteorites” of this article in order to reflect the most up-to-date “trendy” topics on meteoritic studies.

Historical Scientific Investigations of Meteorites

Meteoroids, meteors, and meteorites are three different names referring respectively to often the same objects before, during, and after entering the Earth’s atmosphere. Meteoroids are debris ejected from asteroidal or planetary bodies by impacts, and most meteoroids (barring those from the Moon) travel in highly eccentric orbits in the solar system. When meteoroids are captured by Earth’s gravity, this interplanetary debris strikes the Earth’s atmosphere and experiences intensive and rapid heating due to the adiabatic compression of the air column in front of the meteoroid. Although most of the meteors we see are caused by small grains from comets (e.g., spectacular meteor showers) rather than from meteoroids, the burning descents of meteoroids often produce similar atmospheric phenomenon of light and sound emissions. Such atmospheric phenomenon is called a meteor event or more often a fireball, which is a meteor that appears brighter than Venus. Most meteoroids have been fully vaporized or ablated to millions of tiny dust particles during this descent stage; however, some relics survive to land on the ground. These survivors are called meteorites. A more scientific definition of meteorites by Rubin and Grossman (2010) can be rephrased here: a meteorite is a natural, solid object larger than 10 µm in size, derived from a celestial body, that was transported by natural means from the body on which it formed to a region outside the dominant gravitational influence of that body and that later collided with a body larger than itself. The current flux of all meteoritic materials received on Earth is estimated between 49,000 to 56,000 tons per year (Esser & Turekian, 1988).

The concept and relationship among meteoroids, meteors, and meteorites were not known until modern times. For most of human history, meteoroids, the source and precursor of meteors and meteorites, were unknown. Meteors (fireballs) and meteorite fall events were recorded frequently in human history. The first written record of a meteor event can be traced back as early as the 16th century bce according to the Bamboo Annals in China, while the earliest reliable document of a meteorite fall is dated to 687 bce (Pankenier et al., 2008). Countless reports of meteor or meteorite fall events (often ambiguous in Chinese texts) can be found in various annals throughout Chinese history; however, no meteorite has been preserved to modern times. Ancient Egyptians also paid special attention to such phenomenal events and frequently recorded “irons from heaven.” Modern excavations confirmed these records and discovered the earliest use of meteoritic iron in tools such as dagger blades or as funerary goods dated as early as 3300 bce (Johnson et al., 2013; Comelli et al., 2016). Similar meteoritic irons were also excavated in native American burial mounds at Havana, Illinois, which was dated back to around 2,336 years ago (Grogan, 1948; Arnold & Libby, 1951; Wasson & Schaudy, 1971; McCoy et al., 2017). In the ancient Greek and Roman worlds, writers including Aristotle and Pliny the Elder reported meteorite falls as long ago as 469–467 bce (D’Orazio, 2007). In their languages, the Greek word for iron (sideros) is related to the Latin word for star (sidera) (Heide & Wlotzka, 1995). This etymological link suggests that the earliest source of iron was meteoritic iron, long before humanity mastered the techniques of smelting iron from ore minerals. People living in these earliest civilizations generally view these meteorite fall events as supernatural, even as omens from the gods. Meteorites were venerated as sacred objects in temples and shrines (Newton, 1897). However, most meteorites found before the 1800s have been lost to history, as few survived to this day, barring a few newly excavated funerary goods.

The first meteorite fall preserved to today is a stony meteorite that landed at Nogata, Japan in 861 ad; though it remained unknown and unstudied until much later in history (Shima et al., 1983). The 1492 fall of the Ensisheim meteorite in France caused a sensation among the general public and raised interest among Renaissance intellectuals, thanks to the recent innovation of the printing press (Marvin, 2006). The name, Ensisheim, is representative of the convention that meteorites are named after the place they fell or were found. The fall of the Ensisheim meteorite was viewed as a sign from God by King Maximilian, and as such one piece was sent to the Vatican. The event inspired some of the contemplation on the origins of meteorites. It was not until 1794, 300 years after the fall of the Ensisheim meteorite, that German physicist Ernst E. F. Chladni proposed the cosmic origin of meteorites and the associated meteor (fireball) events in his book titled Über den Ursprung der von Pallas gefundenen und anderer ihr ähnlicher Eisenmassen und über einige damit in Verbindung stehende Naturerscheinungen (On the Origin of the Mass of Iron Found by Pallas and of Other Similar Ironmasses, and on a Few Natural Phenomena Connected Therewith) (Marvin, 1996). Before Chladni’s book popularized the idea of the cosmic origin of meteorites, the reports of falling stones and iron were largely regarded as unbelievable “folk tales” told by ignorant villagers, or “scientifically” explained as volcanic ejecta from the Earth or even from the Moon (Marvin, 1996). The concept of a cosmic origin was previously expressed by the ancient Greek philosopher Diogenes of Apollonia and even Chladni’s contemporaries such as G. C. Lichtenberg, whom Chladni acknowledged. Nevertheless, Chladni was the first to systematically explore and actively advocate his theory. He is rightfully credited as founding a new branch of modern science, meteoritics, or the science of studying meteorites.

Although Chladni’s 1794 book popularized the hypothesis of the cosmic origin of meteorites, this idea initially received widespread rejection from his contemporary natural scientists, and it took many years before it was largely accepted. Fortunately, several instances of meteorite falls and increasing analytical abilities occurred in the following decade (Marvin, 1996) with new witnessed falls at Siena (1794, Italy), Portugal (1796), Mulletiwu (1795, Sri Lanka), Salles (1798, France), Benares (1798, India), L’Aigle (1803, France), and Weston (1807, United States), as well as the chemical analysis of many of these new meteorites by Edward C. Howard and Jacques-Louis de Bournon (1802). Only after the mounting evidence became so convincingly supportive of Chladni’s hypothesis, in part due to the discovery of dozens of asteroids in the early 19th century, did the cosmic nature of meteorites finally become recognized. Since then, the interest in collecting meteorites gradually increased.

Sources of Meteorites

The number of meteorites found and accumulated by private collectors and public institutions has increased exponentially compared to the handfuls at the beginning of the 18th century. Major European natural history museums in Vienna, Berlin, Moscow, Paris, and London all established their sizable meteorite collections around this period (McCall et al., 2006). Many avid private collectors were also successful in hunting and trading meteorites; for example, Harvey H. Nininger (1887–1986) built and curated the largest personal collection of meteorites in his time. The first time all known meteorites were counted and listed was in the publication of the first edition of the Catalogue of Meteorites, published in 1923 by the British Museum. A total of 1,160 meteorites were documented (Prior, 1923). The first meteoritic research organization, the Meteoritical Society, was founded in 1933 to promote the discovery and study of meteorites. In 1953, the first scientific journal Meteoritics (now Meteoritics & Planetary Science) was created by the Meteoritical Society. In 1964, 40 years after the first edition of the Catalogue of Meteorites, the number of known meteorites had increased to 1,548 (Heide, 1964).

Surprisingly, the number of known meteorites exploded in the latter half of the 20th century. As of January 29, 2019 there are 60,155 approved meteorites in the Meteoritical Bulletin Database. This upsurge is not due to any increase in the frequency of meteorite falls providing more samples, nor entirely due to the rapid rise of the economic value of meteorites that might motivate more intensive searches. Although the Space Race between the United States and the Soviet Union and the return of samples from the Moon in 1969 intensified and inspired the research of planetary sciences including meteoritics, the collection of so many meteorites in the late 20th century was due to some unexpected discoveries. Two large meteorite source regions that had been previously inaccessible were identified: the cold desert (Antarctica) and hot deserts such as the Sahara in the northwest Africa. The first meteorite was collected from Antarctica in 1912 during Sir Douglas Mawson’s Australasian Antarctic Expedition (Bayly & Stillwell, 1923). Only three more Antarctic meteorites were found between 1912 and 1969. The mass discovery of Antarctic meteorites began in 1969 with the recovery of nine meteorites from eastern Antarctica’s Yamato Mountains by the Japanese Antarctic Research Expedition (Yoshida et al., 1971; Yoshida, 2010). Antarctica has since become the region of origin for the majority of meteorites (see Figure 1). Government-funded expeditions to Antarctica in the past 40 years by Japan, the United States, Europe, and China have collected 65% of all known meteorites. Although in theory, meteorites can randomly land on any locality on Earth, Antarctica is the ideal place on Earth to “hunt” meteorites because (1) dark-colored meteorites can be readily recognized against the austere background of ice; (2) the base rock of Antarctic continent is covered with thick ice, so few Earth rocks are exposed on the surface to be mixed with meteorites; (3) the low temperature and low humidity preserve meteorites by limiting their weathering and erosion; (4) Antarctica was not visited or inhabited by humans until the last century; thus its remote regions (especially blue ice fields where meteorites are found) were not disturbed; and (5) most importantly, there is a natural “concentration” system in Antarctica: Ice sheet flows, which carry meteorites as they slowly move, stop at the Transantarctic mountains. Meteorites accumulate at the feet of these mountains and become exposed after intense katabatic winds cause ice sublimation. This process works like a “conveyer belt” collecting meteorites over a vast region and depositing them into a confined area (Cassidy et al., 1977; Cassidy & Rancitelli, 1982).

Twenty-five percent of all meteorites have been found in hot desert regions such as the Sahara desert and Arabian Peninsula (primarily in the Sultanate of Oman). Unlike the Antarctic meteorites retrieved due to public-funded efforts, hot desert meteorites are predominantly found by nomads and private meteorite dealers. Hot desert meteorites are more prone to terrestrial weathering and contamination compared to Antarctic meteorites (Crozaz & Wadhwa, 2001; Al-Kathiri et al., 2005; Jull, 2006). Beyond Antarctica, the Sahara desert, and the Arabian Peninsula, only 10% of known meteorites have been collected from other parts of the Earth (see Figure 1). Although more than 60,000 individual meteorites have been documented, many of them belong to the same meteorite falls as they fragmented during the descent or impact with the ground. For example, during the 1947 Sikhote-Alin (Russia) iron meteorite shower, 8,282 various-sized specimens of the same meteorite have been collected and the total mass is about 23 tons (Krinov, 1966). In many cases when the falls were not witnessed, the fragments of the same meteorites may be found at different times or non-adjacent locations. Thus they are cataloged with different names and numbers. Only some of them can be linked to each other when scientists are confident that they are paired and from the same fall event. Therefore the exact number of meteorite falls that produced these individual pieces of meteorites is not easy to estimate.

The mass of a given meteorite can also vary by several orders of magnitude. The five largest single-piece meteorites (rather than the total mass of a single meteorite fall), weigh 60 tons (Hoba, Namibia), 37 tons and 31 tons (two pieces of Campo del Cielo, Argentina), 31 tons (Cape York, Greenland), and 30 tons (Aletai, China). In contrast, the average mass of Antarctica meteorites is only 154 grams, and some are smaller than 1 gram (down to a few milligrams). Most meteorites collected in Antarctica are relatively small in size because meteorites are so much easier to recognize in Antarctica compared to most other places on the Earth, where such rocks are difficult to distinguish. So, while Antarctica is the place from which 65% of all meteorites have been retrieved, by mass that number dwindles to merely 0.8%. Figure 1 shows the geographic distribution of meteorites calculated by both number and mass. On continents populated by people for thousands of years and dominated by terrestrial rocks, only relatively large meteorites are found, and the total number is much less than that from Antarctica. The geographic distribution of known meteorites is severely biased due to many factors such as accessibility and preservation conditions, and this location distribution does not reflect the true distribution of meteorite falls.

Few meteorites are observed during their descents; only 2% of meteorites are “falls,” while 98% are “finds,” with no witnessed meteorite fall events associated with them. Meteorite “falls” are recovered shortly after landing, so they are fresh samples with minimal terrestrial contamination and weathering alteration. In contrast, meteorite “finds” may have been exposed and survived on the surface of the Earth for thousands to hundreds of thousands of years (Jull, 2006). Meteorites that landed in cold and dry environments (i.e., Antarctica) generally survive for longer periods than those in warm and humid environments.

MeteoritesClick to view larger

Figure 1. The proportions of meteorites (by number and by mass) found in different regions on Earth. Data from the Meteoritical Bulletin Database as of January 29, 2019.

Meteorite Classification and Origin

Traditionally, meteorites are first classified into stony, iron, or stony-iron meteorites based on whether they are primarily composed of rocks or metals, or a roughly equal amount of both. The appearances of stony, iron, and stony-iron meteorites are sharply different; so this classification scheme can be used without the assistance of microscopes or chemical analysis. By number, about 95% of all meteorites are stony meteorites, 4% are iron meteorites, and 1% are stony-iron meteorites. The five largest meteorites, mentioned above, are all irons. Stony meteorites are further divided into two categories. After removing the fusion crusts (thin, glassy coatings formed by extreme heating during entry of Earth’s atmosphere), a large proportion of stony meteorites contain millimeter-sized spherules (see Figure 2). This is a unique texture that does not appear in rocks formed on Earth. The stony meteorites bearing this unique spherule texture were recognized early in the history of meteoritics and named chondrites by Gustav Rose in 1864, from the Greek word meaning “grains or seeds” (Rose, 1864; Connolly & Desch, 2004). The millimeter-sized spherules within chondrites are thus called chondrules. Stony meteorites that do not exhibit any chondritic textures (i.e., without any chondrules) are called achondrites. Amongst stony meteorites, chondrites are far more common than achondrites. Chondrites comprise 94.1% of all stony meteorites, with achondrites as the remaining 5.9%.

MeteoritesClick to view larger

Figure 2. Photos of sawn faces of two chondrites showing millimeter-sized spherules (chondrules) inside. A: Murchison (CM2). B: Allende (CV3). The unit of the scale is 1 mm, and images are 20 mm in width.

Photo Credit: Randy Korotev.

After 200 years, this simple division into stony, iron, and stony-iron, the classification scheme has evolved into a complicated system with many groups and clans (Krot et al., 2014) that consider textures, mineralogy, bulk elemental and isotopic compositions, as shown in Figure 3. The modern classification scheme aims to link meteorites that formed in similar ways, to elucidate their origins and identify those with common parent bodies. It first divides meteorites into two major categories: undifferentiated meteorites (chondrites) and differentiated meteorites (achondrites, iron, and stony-iron) mainly based on whether they contain chondrules or not (with rare exception such as CI chondrites). Thus these two categories (undifferentiated and differentiated) can be synonymously called chondritic and non-chondritic. The bulk composition of chondrites resembles that of the Sun, and they are considered some of the most primitive (unaltered) and oldest materials in the solar system (Scott & Krot, 2014). Chondrites originate from undifferentiated parent bodies (e.g., asteroid 25143 Itokawa), where no extensive heating occurs, and thus iron metal and rocky material are not separated to form mantles and cores. In contrast, differentiated meteorites (achondrites, iron, and stony-iron) are from differentiated planets (e.g., Mars) or differentiated asteroids (e.g., asteroid 4 Vesta). Because of the residual heat from accretion and the heating from radioactive isotopes’ decay, differentiated planetary bodies experienced global-scale melting and separation of immiscible liquid iron and rocky silicates. The metal sinks to the center of the bodies, forming a planetary core. The residual silicates form the mantles and crusts. This process is called planetary differentiation.

Generally speaking, differentiated stony meteorites (achondrites) come from the mantles and crusts of differentiated bodies. Iron meteorites are from fragments of cores of destroyed differentiated bodies. Stony-iron meteorites possibly come from the mixing of the core metal and crustal/mantel silicates during asteroidal break-up (Benedix et al., 2014).

One intriguing class of meteorites that do not fit this classification scheme is the primitive achondrites class. They do not contain any chondrules (i.e., chondritic texture; however some have relict chondrules) and have metamorphic (or igneous) textures., Primitive achondrites, however, have similar bulk compositions to chondrites, which indicate incomplete separation of iron and silicates in the parent-bodies. They probably represent samples from asteroids that experienced a higher degree of heating and partial melting on chondrite parent bodies (Mittlefehldt, 2014a). In the current classification scheme, primitive achondrites are categorized as differentiated meteorites but bear a close chemical relationship with chondrites (see Figure 3). The taxonomy of chondrites, achondrites, iron meteorites, and stony-iron meteorites will be discussed in detail in the following subsections. However, other than the groups and clans mentioned below, there are also a growing number of ungrouped meteorites, samples that are unique (or less than five) and do not belong to any established groups. As more meteorites are discovered, it is expected that new groups will be established in the future (Krot et al., 2014).

MeteoritesClick to view larger

Figure 3. The classification of meteorites based on Krot et al. (2014).

Chondrites

Chondrites are conglomerate rocks that contain various components: (chondrules, refractory inclusions, Fe-Ni metals, and fine-grain matrix. Chondrites are like “sedimentary” rocks, which accreted from chondritic components that formed at different times and localities in the solar protoplanetary disk. Chondrites are considered the most primitive materials of the solar system that are available to science, as (1) the refractory inclusions in chondrites represent the oldest solids ever analyzed (4567.30 ± 0.16 Ma; Amelin, 2002; Amelin et al., 2010; Connelly et al., 2017; Connelly et al., 2012); (2) the bulk chemical compositions of chondrites, especially the CI group, resemble those of the Sun (Lodders, 2003); and (3) chondrites contain presolar grains that formed from other stars or supernovae that predate the formation of the solar system (Zinner, 2014). One needs to be cautious here: Although chondrites are primitive and some components are the oldest in the solar system, it does not mean that chondrite parent bodies were formed before differentiated parent bodies. The ages of chondrites are not necessarily older than those of differentiated meteorites (see detailed discussion on chronology in the section “Achondrites”).

Chondrites are divided into three classes: carbonaceous, ordinary, and enstatite chondrites. These three chondrite classes can be further categorized into 13 chondrite groups (see Figures 3 and 4): eight groups of carbonaceous chondrites (CI, CM, CO, CR, CV, CK, CB, and CH); three groups of ordinary chondrites (LL, L, and H); and two groups of enstatite chondrites (EL and EH). In addition, there are two groups of chondrites, the R and K groups, that do not belong to carbonaceous, ordinary, or enstatite chondrite classes (Krot et al., 2014). The name “carbonaceous” chondrites implies that they are carbon-rich. Some carbonaceous chondrites—CI, CM, and CR—do contain significant amounts of carbon (2% to 5%) (Pearson et al., 2006), but others have carbon contents of 1% or less.

The differences between carbonaceous and non-carbonaceous chondrites mostly depend on the following three aspects (Krot et al., 2014): (1) bulk chemistry (i.e., refractory and volatile elements abundances and patterns); (2) mass-independent isotopic compositions (i.e., Δ‎17O, ε‎48Ca, ε‎50Ti, ε‎54Cr, ε‎64Ni, ε‎92Mo, ε‎100Ru and μ‎142Nd; see Figures 5 and 6); and (3) modal abundances of chondritic components (i.e., refractory inclusions, and matrix/chondrule ratio). With a few exceptions, carbonaceous chondrites are high in refractory elements (>0.95 when normalized to the Si abundance in CI), lighter oxygen isotope (Δ‎17O ≤ −2 ‰), refractory inclusion abundance (≥0.1vol%), and matrix relative to chondrule abundance ratio (≥0.9). In contrast, non-carbonaceous chondrites (ordinary and enstatite chondrites plus K and R groups) are low in refractory elements (≤0.95 CI abundances when normalized to Si), lighter oxygen isotope (Δ‎17O ≥ −1 ‰), refractory inclusion abundances (≤0.1vol%) and matrix relative to chondrule abundance ratio (≤0.9). These differences between carbonaceous chondrites and non-carbonaceous chondrites, especially the mass-independent isotopic dichotomy (Figure 6), suggest that carbonaceous chondrites and non-carbonaceous chondrites possibly formed at physically separated (by Jupiter) regions of the solar system (Warren, 2011; Kruijer et al., 2017; Scott et al., 2018).

For the eight groups of carbonaceous chondrites (CI, CM, CB, CO, CV, CK, CR, and CH), their group names comprise the letter C (indicating carbonaceous) plus the initial letter of a typical meteorite name within the group: CI (Ivuna-like), CM (Mighei-like), CO (Ornans-like), CR (Renazzo-like), CH (ALH 85085-like; also means “high” in metal), CB (Bencubbin-like), CV (Vigarano-like), and CK (Karoonda-like). As hinted in the name, ordinary chondrites are the most common type of chondrites, stony meteorites, or meteorites in general (see Figure 4). For the three groups of ordinary chondrites (LL, L, and H), their group names indicate the abundance of total bulk Fe (in either Fe metal, sulfide, or silicates). H ordinary chondrites contain “high” abundance of total Fe (~27 wt.%), while L and LL ordinary chondrites have “low” abundance of total Fe (~22 wt.%). Relative to L chondrites, LL chondrites are not only “low” in total Fe contents, but also “low” in Fe metal and FeS abundances. Enstatite chondrites are named because enstatite (MgSiO3), a pyroxene, is their dominant silicate mineral, rather than olivine [(Mg,Fe)2SiO4] as in ordinary chondrites. For the two groups of enstatite chondrites (EL and EH), EH chondrites contain more metallic iron than EL chondrites. Enstatite chondrites have nearly identical mass-independent isotopic compositions (e.g., Δ‎17O, ε‎50Ti, and ε‎54Cr; see Figure 6) to those of the Earth and Moon (Clayton et al., 1984); thus it has been proposed that enstatite-chondrite-like materials were the building blocks of the Earth (Javoy, 1995; Lodders, 2000; Dauphas, 2017). R (Rumuruti-like) and K (Kakangari-like) groups are significantly different from carbonaceous chondrites in terms of refractory element abundances and are different from ordinary chondrites in terms of matrix/chondrule modal abundance ratios. K and R also have unique mass-independent isotopic compositions (i.e., Δ‎17O).

MeteoritesClick to view larger

Figure 4. The proportions of each type of stony meteorites (by number). Data from the Meteoritical Bulletin Database as of January 29, 2019.

Petrologic type: A secondary classification scheme first defined by van Schmus and Wood (1967) has been widely used to describe the textural and mineralogical properties of chondrites. The scales vary from 1 to 6, with 3 as the most primitive (unequilibrated) type. From type 2 to 1, the degree of low-temperature aqueous alteration increases; while from type 4 to 6, the degree of thermal metamorphism escalates. Van Schmus and Wood (1967) implied that types 2–6 formed from type 1 chondrites, but McSween (1979) corrected this and showed that type 3 chondrites were the starting materials. Chondrule and matrix boundaries are sharply defined in types 3 and 2; however, the boundaries become poorly defined with increasing degrees of thermal metamorphism. For type 1 (mostly CI) chondrites, all chondrules have been destroyed due to strong aqueous alteration. The carbon and water contents are high in types 1 and 2 due to aqueous processes on parent bodies. Table 1 shows the petrologic type and percentages for each chondrite group. Generally speaking, CI, CM, and CR chondrites are all aqueous altered (types 1 and 2), while other types of chondrites are more inclined to thermal alterations (types 4 to 6). Unequilibrated type 3 ordinary chondrites (and CV and CO chondrites in some cases) can be further assigned to subtypes between 3.0 and 3.9 based on their sensitivity to induced thermoluminescence (Sears et al., 1980).

Table 1. Petrologic Types of Chondrite Groups after van Schmus and Wood (1967); Scott and Krot (2014)

Types

1

2

3

4

5

6

Total no.

Aqueous alteration

Pristine

Thermal metamorphism

No chondrules

Chondrules sharply defined

Chondrules sharply defined

Chondrules well defined

Chondrules readily delineated

Chondrules poorly defined

Carbon (wt.%)

3−5

0.8−2.6

<1.5

<1.5

<1.5

<1.5

Water (wt.%)

18−22

2−16

0.3−3

<1.5

<1.5

<1.5

CI

□□□□

5

CM

□□

□□□

48

CO

□□□□

31

CR

□□

□□□

7%

15

CV

3%

□□□

36

CK

□□

□□□

□□□

□□

□□

CB

□□□□

5

CH

□□□□

7

EL

□□□

□□

□□□

□□

EH

□□□

□□□

□□□

□□

□□

LL

□□

□□

□□□

□□□

□□□

L

□□

□□

□□□

□□□

□□□□

H

3□

□□□

□□□

□□□

□□□□

R

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Shock metamorphism stage: In addition to the petrologic type, the degree of shock metamorphism is a tertiary classification scheme for ordinary and enstatite chondrites (Stöffler et al., 1991; Rubin et al., 1997). This scheme varies from S1 (unshocked) to S6 (strongly shocked) based on the shock effects observed in minerals such as olivine, pyroxene, and plagioclase (solid solution of NaAlSi3O8 and CaAl2Si2O8). The corresponding shock pressures vary from less than 4 GPa (S1) to 75−90 GPa (S6). Higher than shock level 6 meteorites melt due to the impact energy. These melted meteorites are called impact-melt or melt breccias meteorites.

The degree of terrestrial weathering: as mentioned above, falls are meteorites observed falling through the atmosphere and are recovered shortly after landing, while finds are those discovered without an observed fall. The “freshness” of meteorites is important to meteorite studies. The possibility of terrestrial contamination and weathering effects needs to be carefully examined for many finds. For meteorite finds, it is useful to indicate semi-quantitatively the degree of terrestrial weathering. Two classification schemes are widely used. Weathering index A, B, or C represent “minor,” “moderate,” or “severe” rustiness of the Fe-Ni metal, respectively. An additional letter “e” represents visible evaporite minerals. This scheme is mainly used for hand specimens of Antarctic meteorites. For thin-section samples of meteorites, the weathering scale W0 (fresh) to W6 (most weathered) is applied based on the oxidation of metals, sulfides, and silicates (Wlotzka, 1993). In general, meteorite finds preserved in cold and dry environments such as Antarctica are less weathered compared to meteorite finds recovered from hot deserts (Crozaz & Wadhwa, 2001; Al-Kathiri et al., 2005; Jull, 2006). Different types of meteorites weather differently even in the same environments, as some types are prone to stronger terrestrial weathering than others. For example, achondrites survive longer than chondrites in hot deserts because the metal in chondrites rusts, causing a volume expansion that in turn breaks the rock apart (Al-Kathiri et al., 2005).

Chondrites in various classes and groups generally contain four major components: chondrules, refractory inclusions, Fe-Ni metals, and fine-grain matrix. The relative abundances of these four components in chondrites are one of the classification criteria. Carbonaceous chondrites contain more matrix than chondrules (CI have <5 vol.% chondrules and consist of 95 vol.% fine-grained matrix), and they also have higher abundance of refractory inclusions (0.1 to 3 vol.%) than ordinary chondrites (<0.1 vol.%). For ordinary and enstatite chondrites, chondrules are the major constituent (60 to 80 vol.%). Chondrule sizes also vary from one group to another (0.15 to 5 mm in diameters). Chondrites in different classes and groups contain different amount of metal, and the abundance can vary from a few percent (for most ordinary and enstatite chondrites) to more than half of the volume (70 vol.% in CH).

Chondrules are millimeter-sized spherical mineral assemblages, which are structures unique in meteorites. Scientists would not predict the existence of chondrules if they did not exist (Connolly & Jones, 2016). Chondrules are thought to have formed by partial or complete melting during a rapid heating event in the protoplanetary disk. The causes of such heating events are still under debate. Possible explanations include interactions with the early active Sun, impacts and collisions between planetary bodies, radiative heating, shock waves, planetesimal bow shocks, and lightning (Sanders & Scott, 2012; Connolly & Jones, 2016). It is worth mentioning that as early as 1877, Henry C. Sorby hypothesized that chondrules “were originally detached glassy globules, like drops of fiery rain” (Sorby, 1877). Olivine (forsterite; Mg2SiO4) and pyroxene (enstatite) are the most common minerals in chondrules. Chondrules also exhibit various textures and can be classified into porphyritic olivine chondrule, porphyritic olivine pyroxene chondrule, granoblastic porphyritic chondrule, barred olivine (with elongated parallel prismatic olivine crystals) chondrule, and radial pyroxene chondrule. The porphyritic chondrules are recognized by the characteristic 120° triple junction where three olivine grains meet. The textures of chondrules are related to the ambient environment and temperature-pressure condition when they formed. For example, laboratory experiments show that the porphyritic chondrules can crystallize from melts with many nuclei (Connolly et al., 1998), and radial pyroxene chondrules can crystallize from a supercooled melt in a levitated environment (Nagashima et al., 2006).

Refractory inclusions are phases that primarily contain minerals condensed at high temperatures (>1300K) in the solar nebula. Refractory inclusions include both calcium-aluminum-rich inclusions (CAIs) and amoeboid olivine aggregates (AOAs). CAIs consist of Ca, Al, and Ti-rich minerals such as corundum (Al2O3), hibonite (CaAl12O19), perovskite (CaTiO3), spinel [(Mg,Fe)Al2O4], melilite (solid solution of gehlenite Ca2Al2SiO7 and akermanite Ca2MgSi2O7), Ti-Al-rich diopside (CaMgSi2O6), and anorthite (CaAl2Si2O8). With the exception of diopside and anorthite, these minerals are all refractory minerals. Corundum, hibonite, perovskite, spinel, and melilite are formed at higher temperatures (e.g., >1300K at the pressure ~10−4 bar) than iron, olivine, and pyroxene in a solar composition gas over a wide pressure-temperature range, while diopside and anorthite condense in about the same temperature range as olivine and pyroxene, and are formed from the earlier condensed refractory minerals (Lord, 1965; Larimer & Anders, 1967; Grossman, 1972; Ebel & Grossman, 2000; Ireland & Fegley, 2000; Petaev & Wood, 2004). Because of the highly refractory characteristics of CAIs, the relative age calculated from the short-lived 26Al radionuclide, and also the absolute age calculated via Pb-Pb dating (4567.30 ± 0.16 Ma; Amelin, 2002; Amelin et al., 2010; Connelly et al., 2017; Connelly et al., 2012), CAIs are considered the earliest solids to condense from the cooling solar nebula. AOAs contain extremely fine olivine grains and also a refractory component including Al-diopside, anorthite, spinel, and melilite.

Chondrites also contain metal, within both chondrules and the matrix. In contrast achondrites have little metal and sulfide because of the depletion of metal and sulfide during planetary differentiation. Metal in chondrites is a Ni-rich iron alloy. They were formed during condensation of the solar nebula with forsterite and enstatite in the temperature range of ~1350−1450K (Campbell et al., 2005). There are also refractory noble metal nuggets containing mostly iridium, osmium, ruthenium, molybdenum, tungsten, and rhenium (Wark & Lovering, 1976). These refractory noble-metal nuggets were formed from the condensation of the solar nebula along with CAI minerals at temperature >1600K (Palme and Wlotzka, 1976; Berg et al., 2009). Refractory noble-metal nuggets are much less abundant than Ni-rich Fe alloy. Most types of chondrites contain a few percent of metal, although CH and CB contain large proportions of metal (>20 vol.%).

The matrix is composed of ultra-fine-grained (10 nm to 5 μ‎m) minerals, mostly Fe-rich olivine that can’t be recognized with an optical microscope. There are also Mg-rich olivine grains, opaque minerals, and amorphous silicates and organics. Matrix also contains recognizable chondrules and CAI fragments. Matrix is a mixture of many types of materials. While some people call it the “garbage can” in chondrites, others view it as the “gold mine.” Matrix has been considered to be a low-temperature product in contrast to chondrules, AOAs, and CAIs that were formed at a higher temperature, but most of the components in chondrite matrices appear to have formed at high temperatures.

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Figure 5. Schematic oxygen “three-isotope” plot of meteorites. Date are from Clayton, (1993); Krot et al. (2014). The delta (δ‎) notations are defined as per mil (‰) deviations of 18O/16O (or 17O/16O) ratios in meteorites relative to those ratios in the Standard Mean Ocean Water (SMOW). Any mass-dependent fractonation should plot on the Terrestrial Fractioantion (TF) line or on lines parallel to the TF line (slope ~0.5). However, the majority of chondrites and some achondrites (e.g., ureilites) fall on a mass-independent fractonation line (CCAM: Carbonaceous Chondrite Anhydrous Mineral; slope ~1). The Δ‎17O values (see Figure 6) are calculated as the deviations from the TF line.

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Figure 6. ε‎54Cr versus ε‎50Ti and Δ‎17O in meteorites. The epsilon (ε‎) and Delta (Δ‎) notations represent the deviations of ratios (54Cr/52Cr, 50Ti/47Ti, 17O/16O) in meteorites relative to those ratios of the Earth (after mass-dependent isotopic fractionation correction). By definition, the ε‎54Cr, ε‎50Ti and Δ‎17O values of the Earth are zeros. Different groups of meteorites have unique mass-independent isotopic compositions. Note that there is a “Warren gap” between carbonaceous chondrites and other meteorites, which is named after Warren (2011). This gap has been explained as that carbonaceous chondrites are formed at outer solar system while others at inner solar system (Warren, 2011; Kruijer et al., 2017; Scott et al., 2018). The data is compiled by Dauphas (2017) from original data (Shukolyukov & Lugmair, 2006; Trinquier et al., 2007; Leya et al., 2008; Trinquier et al., 2009; Qin et al., 2010; Yamakawa et al., 2010; Zhang et al., 2012; Schiller et al., 2014; Göpel et al., 2015; Young et al., 2016).

Achondrites

Differentiated meteorites exhibit textures indicating recrystallization after partial or complete melting; they are igneous rocks or breccias composed of igneous rocks. They originate from asteroids or planets that have chemically differentiated into core, mantle, and crust. Achondrites are representative of the mantle and crust, while iron and stony-iron meteorites are possibly from the core and core-mantle boundary.

Achondrites include primitive achondrites and differentiated achondrites from asteroids (e.g., 4 Vesta) and planets (e.g., Mars). The classification and proportions of each group are shown in Figures 3 and 4. Achondrites are different from chondrites in many aspects: (1) achondrites are chemically fractionated (different abundances and patterns of refractory/volatile and lithophile/siderophile elements); (2) they contain sparse metal and sulfide; (3) they do not have the textures of chondrites but instead show igneous or metamorphic textures; (4) they do not contain chondrules, CAIs, AOAs, or other chondritic components; and (5) they have different signatures of mass-independent isotope compositions (i.e., Δ‎17O, ε‎48Ca, ε‎50Ti, and ε‎54Cr; see Figure 6). Enstatite achondrites are an exception because they are identical to enstatite chondrites in mass-independent isotope compositions.

Primitive Achondrites

Primitive achondrites are those meteorites having bulk compositions similar to chondrites but which exhibit igneous textures. Primitive achondrites include three subgroups: the acapulcoite-lodranite clan, winonaites, and brachinites (Mittlefehldt, 2014a). They are probably formed by ultra-metamorphism of chondrites, or they represent the residues of very-low-degree partial melting (Krot et al., 2014).

Acapulcoite-lodranite clan is the prototype of primitive achondrites, and they have a similar mineral assemblage to that of ordinary chondrites: orthopyroxene, olivine, diopside, plagioclase, Fe-Ni metal, and sulfide. The bulk chemical compositions of acapulcoite-lodranite meteorites are similar to those of ordinary chondrites, too (especially of the H group). Unlike ordinary chondrites, acapulcoite-lodranite clan meteorites do not have any chondrules, and they are fine- to medium-grained equigranular mafic igneous rocks (Mittlefehldt et al., 1998). The oxygen isotopes of acapulcoite-lodranite meteorites are also clearly different from those of ordinary chondrites (see Figure 6). Acapulcoite-lodranite meteorites have a relatively high abundance (~10−20%) of metal compared to differentiated achondrites, which have lost metal to core formation in their parent bodies. The formation mechanisms of acapulcoite-lodranite meteorites are proposed to be the result of the extensive high-temperature metamorphism and anatexis of primitive, reduced chondritic material (Mittlefehldt, 2014a).

Winonaites are also chondritic in term of bulk chemistry, mineral assemblage, and modal abundances (olivine, orthopyroxene, clinopyroxene, plagioclase, troilite, and metal). They do not have chondrules (except relicts of chondrules in rare occasions), and they show equigranular igneous textures. All these characteristics of winonaites are similar to those of the acapulcoite-lodranite clan (Krot et al., 2014); however, they bear distinct oxygen isotope compositions. The oxygen isotope compositions also show links between winonaites and silicate inclusions from main-group IAB irons (Clayton & Mayeda, 1996). They can be combined to be called the winonaite-IAB-iron silicate inclusion clan. This genetic relationship appears to require that both primitive meteorites and differentiated iron meteorites originate from the same parent body. The formation mechanism of such winonaite-IAB-iron parent body is thus still under debate. One possible explanation is that this clan was formed from a chondritic parent body where differentiation was ongoing but was interrupted by a catastrophic impact. Materials from various depths (partially formed metallic core, silicates, and highly metamorphosed chondritic materials on the top) were mixed (Benedix et al., 2000).

Brachinites are a small group of olivine-rich ultramafic primitive achondrites. Their type specimen, Brachina, was found in South Australia in 1974 and it was shown to be chemically similar to chondrites but has undergone some segregation of sulfide melt and exhibits equigranular igneous textures (Nehru et al., 1983). Brachinites contain olivine (79–95 vol.%), augite (3–15 vol.%), plagioclase (0–10 vol.%), and trace amounts of orthopyroxene, chromite, phosphates, Fe sulfides, and Fe-Ni metal (Krot et al., 2014). Olivines are medium- to coarse-grained, and they show triple-junction equigranular texture; while augite and plagioclase are fine-grained. The bulk chemistry of brachinites is near chondritic in term of lithophile and siderophile elements; thus they were considered primitive achondrites. Some brachinites are ultra-metamorphosed chondrites and others are residues of low-degree partial melting (Nehru et al., 1983; Nehru et al., 1992, 1996; Day et al., 2012). Alternatively, brachinites were also proposed to be cumulates due to the coarse-grained and oriented cumulate textures (Warren & Kallemeyn, 1989; Swindle et al., 1998; Mittlefehldt et al., 2003).

Achondrites from Asteroids

The majority of achondrites are from asteroids, and most come from a single asteroid—4 Vesta (see Figure 4). The second largest asteroid in the asteroid belt, 4 Vesta was discovered in 1807 by German astronomer H. W. Olbers. In 1864, eucrites, the most common type of stony meteorites after chondrites, were recognized and discussed by German mineralogist G. Rose (Rose, 1864; Tschermak, 1885; Drake, 2001). However, until 1970, these eucrites were not linked to 4 Vesta. Based on spectroscopic observations, the reflectance spectra of 4 Vesta are characteristically similar to those of the eucrites (McCord et al., 1970). This link has been confirmed by the recent NASA Dawn close-flyby mission to 4 Vesta and 1 Ceres (McSween et al., 2013). Eucrites and associated meteorites form the HED (Howardite-Eucrite-Diogenite) clan because of their identical oxygen isotope composition (Greenwood et al., 2005). Howardites are polymict breccias of mixtures of different proportions of eucrite and diogenite clasts. Eucrites are basalts (non-cumulates) or cumulate gabbros. Basaltic eucrites contain pigeonite (low-Ca clinopyroxene), plagioclase, and a minor amount of silica, ilmenite, and chromite (Duke & Silver, 1967). Basaltic eucrites can be subdivided based on incompatible trace elements: “Main Group,” “Nuevo Laredo Trend,” and “Stannern Trend” (Stolper, 1977; Consolmagno & Drake, 1977; Reid & Barnard, 1979; Yamaguchi et al., 2009). Cumulate eucrites have mineral assemblages similar to basaltic eucrites, but they display cumulate textures, and their pyroxenes are Mg-rich. Most eucrites are brecciated and could be monomict (one lithology) or polymict (containing multiple lithologies). Diogenites are orthopyroxenites, consisting chiefly of ~90 vol% coarse-grained orthopyroxene, and accessory minerals including olivine, chromite (FeCr2O4), troilite (FeS), and metal (Mittlefehldt et al., 1998). Most diogenites are also brecciated. In term of bulk chemistry, HEDs are highly fractionated compared to chondrites, and are significantly depleted in moderately and highly volatile elements as well as moderately and highly siderophile elements (Mittlefehldt, 1987; Warren et al., 2009; Dale et al., 2012). Vesta was formed early in the solar system, and the ages of HED meteorites are as old as 2–10 Ma after the formation of CAIs (Trinquier et al., 2008; Iizuka et al., 2014; Iizuka et al., 2015).

In 2008, a small (~4m) asteroid (2008 TC3) hit the Earth, and around 700 pieces of this asteroid were recovered from the predicted landing region (Jenniskens et al., 2009; Goodrich et al., 2015). This event was the first time that an asteroid was discovered in space, tracked while descending through the atmosphere, and immediately recovered. The collected meteorite, named Almahata Sitta (Sudan; “station 6” in Arabic), is a polymict ureilite. Ureilites are coarse-grained ultramafic rocks dominated by olivine and pyroxene, with minor phases such as carbon, metal, sulfide, and phosphides (Goodrich, 1992; Mittlefehldt et al., 1998). They are the second largest group of achondrite (see Figure 4). The existence of carbon phases (up to 5%; mainly graphite) and metallic iron in the matrix of ureilites suggest that the ureilites have been formed in an extremely reducing environment (Mittlefehldt et al., 1998). Since 1888, nano-diamonds have been found in the type specimen Novo-Urei and other ureilites (Jerofejeff & Latschinoff, 1888; Ringwood, 1960). Ureilites experience various degrees of impact shocking, from unshocked to up to 100 GPa (Mittlefehldt, 2014a). The oxygen isotopes of ureilites are unique compared to other groups of achondrites. Unlike other achondrites, ureilites do not have homogeneous Δ‎17O within the group, and their oxygen isotopes do not fall on a 0.5-slope mass-dependent fractionation line but on a 1-slope carbonaceous chondrite anhydrous mineral line (Clayton & Mayeda, 1996). Most ureilites are coarse-grained olivine-pigeonite ureilites: They are the partial melting residue from a carbon-rich chondritic parent body (Warren & Kallemeyn, 1992; Scott et al., 1993; Goodrich et al., 2002), though no corresponding complementary rock types are recognized. Other ureilites are olivine-orthopyroxene-(augite) ureilites. They are poikilitic in texture and are thought to have formed as igneous cumulates (Goodrich et al., 2001).

Aubrites are also called enstatite achondrites because they have identical enstatite-dominated mineralogy and Δ‎17O composition as enstatite chondrites, although they are not from the same parent body. Aubrites are achondrites, as they do not have any chondrules and are depleted in siderophile elements, in contrast to enstatite chondrites. Aubrites are FeO-poor enstatite orthopyroxenites, containing ~75–95 vol.% FeO-free enstatite, Si-bearing kamacite, and high abundances of rare sulfide minerals (troilite, oldhamite, daubreelite, and alabandite). They formed under extremely reducing conditions similar to enstatite chondrites (Keil, 2010; Mittlefehldt, 2014a). All the aubrites (except Shallowater) are brecciated and appear to have formed on the same parent body. In contrast, the Shallowater aubrite is the only unbrecciated aubrite and likely comes from a distinct parent body (Keil, 2010).

Angrites are amongst the oldest basaltic rocks in the solar system, and are the most alkali-depleted rocks (Keil, 2012). Angrites are a small group of meteorites, with only 28 recognized so far. The name came from the type specimen, Angra dos Reis (Brazil, 1869), which is the only observed fall among the angrites (Keil, 1977). All other angrites are finds from Antarctica, Northwest Africa or Argentina. Angrites comprise Al-Ti-diopside, Ca-rich olivine, kirschsteinite, plagioclase, and minor minerals such as spinel, troilite, whitlockite, ulvöspinel-magnetite, and metal (Mittlefehldt, 2014a). They are medium to coarse-grained, and are unbrecciated and unshocked (Keil, 2012). The crystallization age of the oldest angrites is 4564.4 ± 0.1 Ma (Amelin, 2008b, 2008a), which suggests that the angrite parent body (and other planetesimals) accreted and differentiated extremely soon (<3 Ma) after formation of the solar system (4567.30 ± 0.16 Ma; Amelin, 2002; Amelin et al., 2010; Connelly et al., 2017; Connelly et al., 2012). It is most likely that angrites come from a differentiated asteroid rather than from a planet-sized body (Keil, 2012).

Planetary Meteorites

Approximately 99.8% of all known meteorites are pieces of asteroids, and only a few rare meteorites come from planetary bodies such as the Moon (0.1%) and Mars (0.1%). It is possible that meteorites were delivered to the Earth from Mercury (Gladman et al., 1996; Gladman & Coffey, 2009) and there are claims of such mercurian meteorites (e.g., Papike et al., 2003), although none have been confirmed. Venusian meteorites are much less likely to be found on Earth due to the large gravity and dense atmosphere of Venus preventing materials reaching escape velocity (Gladman et al., 1996). Compared to the meteorites from asteroids, meteorites from planets are much younger due to the more extended magmatic activity on big planetary bodies. In contrast, smaller bodies such as asteroids cooled faster, and there is no lasting magmatic activity after ~4.4 Ga.

Martian meteorites are also called SNC (Shergottites, Nakhlites, and Chassignites) meteorites. The first meteorite belonging to this group is Shergotty, a 1865 fall in India that Gustav Tschermak first studied in 1872 (McSween & Treiman, 1998). However it was not until more than 100 years later that Shergotty and other SNC meteorites were recognized to originate from Mars. This attribution is based upon (1) their young crystallization ages, and more importantly (2) their similar nitrogen and noble gas isotope compositions to those of Mars’ atmosphere measured by the Viking spacecraft (Bogard et al., 2001). Shergottites are the most common type of martian meteorites, and they could be either diabasic, gabbroic, pigeonite-phyric, olivine-phyric, or poikilitic in texture (Walton et al., 2012). Nakhlites are clinopyroxenites or wehrlites. They are coarse-grained and mainly consist of augite and olivine. Chassignites are dunites. Other than shergottites, nakhlites, and chassignites, the fourth type of martian meteorites is an orthopyroxenite ALH 84001. It is the only known martian orthopyroxenite. ALH 84001 was claimed to have “nanofossils” (McKay et al., 1996); however, the biogenic origin is highly controversial. Lherzolitic shergottites, nakhlites, chassignites, and ALH 84001 are all cumulates. Most martian meteorites have young crystallization ages (e.g., 100–500 Ma for shergottites; ~1.3 Ga for nakhlites and chassignites; see summary in McSween & McLennan, 2014). In contrast, ALH 84001 represents the ancient martian crust and crystallized at 4.1 Ga (Barrat & Bollinger, 2010).

Lunar meteorites, or lunaites, are rocks found on Earth that were ejected from the Moon by the impact of another body (Korotev, 2005). Meteoroids strike the Moon every day. Any rock on the lunar surface that is accelerated by the impact of a meteoroid to or beyond lunar escape velocity will leave the Moon’s gravitational influence. Most rocks ejected from the Moon become captured by the gravitational field of either the Earth or the Sun and go into orbits around these bodies. Over a period of a few to tens of thousands of years, those orbiting the Earth eventually fall to Earth. Those in orbit around the Sun may also finally strike the Earth tens of millions of years after they were launched from the Moon. The first lunar meteorite was not recognized until 1982 (Bogard, 1983; Warren et al., 1983). More than 350 lunar meteorites have been found today. Intriguingly, all lunar meteorites are finds from Antarctic or “hot” deserts (Sahara and Oman); there are no observed lunar meteorite falls (Korotev, 2012; Korotev and Zeigler, 2014). In addition to lunar meteorites, six U.S. Apollo and three Soviet Luna programs have collected and brought back 382 kg of lunar rock and soil samples. The mass of all known lunar meteorites is about 74% of the mass of the rocks in the Apollo lunar sample collection. Compared to lunar rocks and soil samples collected from the nine landing sites of the Apollo and Luna programs, lunar meteorites represent an almost-random sampling of the Moon (Gladman et al., 1995; Le Feuvre & Wieczorek, 2008; Gallant et al., 2009) and provide us with a better estimate of the composition and mineralogy of the lunar surface than provided by the Apollo and Luna samples (Korotev et al., 2003). Lunar meteorites (and Apollo lunar samples) are classified into two broad categories. Rocks from the lunar highlands are rich in aluminum because they contain a high proportion (60–99%) of plagioclase (anorthite), and are called anorthosite (see Figure 7). Rocks from the maria are basalts that are rich in iron. They are crystalline, igneous rocks consisting mainly of pyroxene, plagioclase, olivine, and ilmenite. All lunar meteorites from the highlands are breccias, composed of fragments of different rocks and are held together by shock compaction or by material that was partially or entirely molten.

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Figure 7. Photo of the sawn face of lunar meteorite Yamato 86032: a breccia with black (impact melt), dark gray, light gray, and white (anorthosite) lithologies. The cube is 10 mm in diameter.

Photo Credit: Randy Korotev.

Iron Meteorites

Iron meteorites are the first extraterrestrial materials that were recognized and utilized by humans thousands years ago. The natural form of metallic iron is extremely rare on the surface of the Earth due to the high oxygen abundance in the atmosphere and hydrosphere, as well as because most of the Earth’s iron was incorporated into the core during the planetary differentiation. Iron meteorites thus appear to be entirely different from any rocks on Earth, and can be easily recognized. Iron meteorites also survive longer on Earth than stony meteorites. It is thus not surprising that the mass proportion of iron meteorites is higher than 85% of all meteorites, and that the 15 largest meteorites are all irons. Compared to the frequency of iron meteorite falls, they are overrepresented in the current meteorite collection due to these biases of preservation and recognition.

Iron meteorites are nearly 100% Fe-Ni alloy, although many contain iron sulfide, phosphide and carbide minerals such as troilite (FeS), schreibersite [(Fe,Ni)3P], and cohenite [(Fe,Ni)3C]. The bulk abundances of Ni in most iron meteorites vary between 5% and 20%, with rare ones containing up to 60% or down to 4% (Mittlefehldt et al., 1998). The low-Ni (<6%) Fe-Ni alloy is called kamacite, while the high-Ni (>30%) alloy is called taenite. Kamacite can (very rarely) be found on Earth, while taenite does not exist on Earth at all. Thus the appearance of taenite and the resulting high-Ni content in iron meteorites are key characteristics to distinguish meteoritic iron from natural or industrial iron on Earth. Another useful characteristic is that industrial metal typically contains more than 100 ppm each of manganese and chromium, while the two elements have a concentration less than 100 ppm in iron meteorites. Many (not all) iron meteorites contain both kamacite and taenite, and the lamellae of kamacite and taenite interlace with each other forming Widmanstätten patterns (see Figure 8). The Widmanstätten pattern is a unique texture that can be only found in iron meteorites (like chondrules in stony meteorites). This intergrowth of kamacite and taenite can be revealed by acid etching (e.g., 2% nitric acid with ethanol) on freshly polished faces. Taenite is more resistant to acids compared to kamacite, so it stands out in positive relief. This characteristic pattern of most common iron meteorites was first discovered by G. Thompson and E. C. Howard and later by Alois von Widmannstätten (Mittlefehldt et al., 1998), who gave the name to this structure. The Widmanstätten pattern can also be shown by burning and oxidizing iron meteorites. Taenite and kamacite have different oxidation rates, and taenite is more resistant to oxidation compared to kamacite. Thus the colors of partially oxidized taenite and kamacite exhibit differently and show Widmanstätten pattern too (Lodders & Fegley, 2011).

The classification of iron meteorites can be based either on structure (class) or bulk chemistry (group) (Buchwald, 1975). Traditionally, iron meteorites can be divided into hexahedrites, octahedrites, and ataxites. Hexahedrites contain only large kamacite crystals without taenite; thus they are low in nickel and have no Widmanstätten pattern. Octahedrites consist of both kamacite and taenite; therefore, they have Widmanstätten pattern and average nickel contents. Octahedrites are the most common class of iron meteorites. Ataxites contain mostly taenite, are high in Ni abundance, and have no Widmanstätten pattern. Ataxites are the rarest class. This structure classification can be further divided according to the bandwidths (coarse, medium, fine and ultrafine) of the kamacite crystals. The bandwidths depend on Ni and P contents and are related to cooling rates, which reflect the sizes of the parent bodies (Mittlefehldt et al., 1998; Yang & Goldstein, 2005). In addition to the classification scheme based on structure, iron meteorites can be classified according to their bulk chemistry. With better analytical ability of siderophile elements (Ni, Ga, Ge, and Ir), four groups (Roman numerals I, II, III, and IV) in the 1950s (Brown & Goldberg, 1949; Goldberg et al., 1957; Lovering et al., 1957) are first subdivided (by adding letters A, B. . .), were later combined (e.g., IAB), and now have finally evolved into 12 different groups (Scott & Wasson, 1975; Krot et al., 2014): IAB, IC, IIAB, IIC, IID, IIE, IIIAB, IIICD, IIIE, IIIF, IVA, and IVB. The reason why Ga and Ge are used in this chemical classification is that Ga and Ge are among the most volatile elements in iron meteorites and there are large variations of Ga and Ge among iron groups due to varying degrees of volatile fractionation prior to solidification (Wasson & Wai, 1976; Wai & Wasson, 1979). Within the same chemical group, Ga and Ge abundances do not significantly vary because they do not fractionate significantly during crystallization of Fe-Ni, unlike Ir, which greatly prefers solid to liquid Fe-Ni (Goldstein et al., 2009). Among all iron meteorites, IIIAB is the most abundant group and comprises one-third of all samples, while IAB is the second largest group, containing ~15% (Krot et al., 2014). There are also ~15% iron meteorites that do not belong to any of the defined groups. They are called ungrouped iron meteorites.

These 12 groups of iron meteorites can be categorized into two “super” groups: magmatic iron meteorite groups and nonmagmatic groups. Magmatic iron meteorite groups include IIAB, IID, IIIAB, IIIE, IIIF, IVA, and IVB, which show evidence of fractional crystallization. They are proposed to have come from cores of differentiated asteroids (Benedix et al., 2014). In contrast, nonmagmatic iron meteorite groups (silicate-rich groups) include IAB, IIE, and IIICD. The variations of Ni and trace elements in these nonmagmatic groups cannot be explained by simple fractional crystallization of metallic cores. They also contain relatively large amount of silicates, which can be linked to primitive achondrites such as winonaites. Nonmagmatic iron meteorite groups might have come from an impact-destroyed partially melted chondritic (undifferentiated) parent body (Benedix et al., 2000), or from impact-melted and mixed parent bodies of ordinary chondrites (Wasson, 2017). The relative formation age of magmatic iron meteorite groups is within 1 Ma of the formation of CAIs (Kleine et al., 2009) and they represent the remnants of the cores of earliest differentiated planetesimals. Nonmagmatic iron meteorite groups are younger, and they were formed at least 5 Ma or more after the formation of CAIs from parent bodies with insufficient short-lived radionuclides (e.g., 26Al) and have not experienced global melting and differentiation (Schulz et al., 2012).

Based on their mass-independent isotopic compositions (i.e., ε‎94Mo and ε‎95Mo) and their relationships to chondrites (Kruijer et al., 2017), iron meteorites can be also classified into two groups: carbonaceous chondrite group (IIC, IID, IIF, IIIF, and IVB) and non-carbonaceous chondrite group (IC, IIAB, IIIAB, IIIE, and IVA). This classification is independent from classification based on bulk elemental chemistry, structure, or texture. “Non-carbonaceous” iron meteorite groups have Mo isotopic compositions falling on the same trend of enstatite and ordinary chondrites, while “carbonaceous” iron meteorite groups are similar to carbonaceous chondrites. Based on Hf-W chronology (Kruijer et al., 2017), “Non-carbonaceous” iron meteorites were formed earlier than “carbonaceous” iron meteorites.

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Figure 8. The sawn, polished, and etched slab of the Gibeon (IVA) iron meteorite showing Widmanstätten pattern.

Photo Credit: Randy Korotev.

Stony-Iron Meteorites

Stony-irons consist of almost equal amounts of nickel-iron alloy and silicate minerals. Stony-iron is the smallest type in the trichotomy classification scheme of meteorites (stony, iron, and stony-iron), and it contains only two groups: pallasites and mesosiderites. Pallasites consist of large and well-defined olivine crystals embedded in an iron-nickel metal matrix. Pallasites are named after German naturalist Peter Simon Pallas, who found the type-specimen in the 1770s and inspired E. F. Chladni’s investigation into meteorites (Lodders & Fegley, 2011). Mesosiderites are breccias containing comparable amounts of fine-grained silicates and Fe-Ni metal. The name is derived from Greek, from “in the middle” (mesos) and “iron” (sideros). Pallasites and mesosiderites are not genetically related, nor do they have a similar composition. They are categorized into stony-iron meteorites simply for convenience.

Pallasites, the better-understood group of stony-irons, are dominated by large olivine crystals surrounded by a continuously enclosing network of Ni-Fe metal. They also contain other minerals such as chromite, low-Ca pyroxene, troilite, schreibersite, and graphite (Krot et al., 2014). Most pallasites are called main-group pallasites. The Eagle Station pallasites grouplet is similar to the main-group pallasites in term of mineral assemblages; however, they are different in mineral chemistry (e.g., Fe content and Fe/Mn ratio in olivines, Ir and Ni abundances in metals) and oxygen isotopes (Mittlefehldt et al., 1998; see Figure 6). The third grouplet of pallasites is pyroxene pallasites wherein pyroxene replaces olivine as the dominant silicate mineral. There are also ungrouped pallasites that do not belong to either main group, Eagle Station or pyroxene pallasites. The olivine in pallasites is consistent with a mantle origin; pallasites are conventionally thought to come from the core-mantle boundary of differentiated asteroids (Mittlefehldt et al., 1998). An alternative explanation is impact mixing of the liquid Fe core from one asteroid and the shallow mantle from another asteroid (Tarduno et al., 2012).

Mesosiderites are even more enigmatic. Although mesosiderites are made of roughly half metal and half silicates (the actual ratios of metal over silicates varies significantly from sample to sample), they appear entirely different from pallasites. The silicates in pallasites are coarse olivine crystals, while mesosiderites consist mainly of fine-grained silicates such as orthopyroxene, olivine, pigeonite, Ca-plagioclase, and silica (Krot et al., 2014). The texture of mesosiderites is also distinct from pallasites. There is a distinct separation between metal and silicate phases in pallasites, while silicates in mesosiderites intermix with the metal alloy. The metal grains are small (mm-sized) and numerous in mesosiderites while metal grains in pallasites are typically cm-sized or larger. All mesosiderites are breccias consisting of lithic clasts of basalts, gabbros, and pyroxenites. Therefore, unlike pallasites, the silicate portions of mesosiderites are more crust-like than mantle-like. The cooling rate of mesosiderites is extremely slow (in fact, the slowest among any natural geological materials) based on the Ni diffusion profiles, which suggest that mesosiderites are cooled within a large (200–400 km in radius) asteroidal body (Haack et al., 1996). The oxygen isotope compositions of mesosiderites overlap with those of HED meteorites (Greenwood, 2006; Greenwood et al., 2005; see Figure 6). However, this genetic relationship between mesosiderites and HED meteorites (asteroid 4 Vesta) has not been confirmed by the NASA Dawn mission, and is still controversial (Mittlefehldt et al., 2012; Scott et al., 2014; Mittlefehldt, 2014b).

Selected Research Themes

Presolar Grains in Meteorites and Stellar Sources of the Solar Nebula

CAIs are the first condensates in the solar system. They have been dated using radionuclides, and define the beginning of the solar system (4567.30 ± 0.16 Ma; Amelin, 2002; Amelin et al., 2010; Connelly et al., 2017; Connelly et al., 2012). There are other grains (minerals or amorphous substance) in chondrites known to predate the formation of the solar system. These grains are called presolar grains (see Figure 9). Due to their extremely small sizes (nanometer to a few micrometers) and other analytical difficulties, there is no direct way to date the ages of presolar grains using radionuclides as is possible with CAIs, chondrules, and bulk meteorites. However, presolar grains are known to be foreign to the solar system as they contain extremely different isotopic compositions from all other objects in the solar system. The isotopic variation among all known solar materials are relatively small (when compared to the large isotopic variation created by stellar sources) and the variation follows the mass-dependent fractionation law (or rarely following mass-independent fractionation; Thiemens & Shaheen, 2014). The extremely different isotopic compositions found in presolar grains cannot have formed within the solar system and must be inherited from the dust grains (“stardust”) that condensed in the stellar outflows of late-type stars or the ejecta of stellar explosion such as supernovae (Zinner, 2014). While most of the stardust that formed our solar system has been destroyed and mixed (to form the solar system), the presolar grains are remnants that survived the evaporation and homogenization processes during the collapse of the interstellar molecular cloud that became the “hot” solar nebula (Lodders & Amari, 2005). Presolar grains were incorporated into the asteroids and comets and are delivered to Earth by the least metamorphosed and least altered chondrites (mainly carbonaceous chondrites, plus some of the least metamorphosed type 3 ordinary and enstatite chondrites), interplanetary dust particles (IDPs), and micrometeorites. Presolar grains have preserved the unique isotopic compositions of their respective parent stars (Zinner, 2014; Floss & Haenecour, 2016).

Classic nucleosynthetic studies have predicted that different nucleosynthetic processes in different stars would produce different isotopic compositions and all these stellar sources mixed together in various proportions contribute to the grand “average” of the solar system (Cameron, 1957; Burbidge et al., 1957). Large isotopic anomalies deviating from this grand “average” had not been observed until 1969. The first isotopic anomalies were found in noble gasses such as neon and xenon (Black & Pepin, 1969; Anders et al., 1973), which suggests that isotopically distinct presolar grains exist in those meteorites. However, those measurements were done on bulk meteorites. The carrier phases of these isotopic anomalies were unknown. They were finally isolated after a tenuous chemical procedure, which has been described as “burning down the haystack to find the needle” (Lewis et al., 1987). Since then, many types of presolar grains such as nanodiamond, silicon carbide (SiC), graphite, silicates (e.g., olivine and pyroxene), oxides (e.g., spinel, corundum, and hibonite), nitrides, and metals have been isolated from the rest of the isotopically normal meteorites or identified in situ within meteorite sections through isotopic imaging with nanometer-scale secondary ion mass spectrometry (NanoSIMS; Zinner, 2014). These presolar grains exhibit huge isotopic differences from the solar system composition in nearly all isotope systems that can be measured in such small grains, such as C, N, O, Al-Mg, Si, Ca, Ti, Fe, and noble gases (Lodders & Amari, 2005). Some grains only show single isotopic anomaly while others show multiple isotopic anomalies in several elements within the same grain.

The abundances of presolar grains also vary in different meteorites. Even in the most primitive meteorites such as carbonaceous chondrites, the abundances of presolar grains are only up to 1,400 ppm (by mass) for nanodiamond, 220 ppm for silicates, 150 ppm for SiC, 80 ppm for oxides, 2 ppm for graphite, and 3 ppb for nitride (Zinner, 2014). For other meteorites containing presolar grains (the least metamorphosed type 3 ordinary and enstatite chondrites), the abundances are even lower. In metamorphosed type 4–6 ordinary and enstatite chondrites, and differentiated meteorites, almost no presolar grains survive (except rarely nanodiamonds).

The extremely different isotopic compositions of presolar grains can help us identify the stellar sources of the solar system by comparisons with astrophysical models, which predict the isotopic compositions of stellar outflows of late-type stars and the ejecta of stellar explosions. Presolar grains likely represent a “biased” sampling of the materials that formed the solar system, as most presolar grains were destroyed, and those that survived are “biased” toward grains (e.g., nanodiamond, the most abundant type of solar grains) that are resilient against destructive processes. However, presolar grains can still provide us a wealth of valuable information that we otherwise have no approach to acquiring. While astronomical observations provide large-scale information on stars and stellar processes, presolar grains represent a direct in situ “snapshot” of the conditions and composition in a single star.

Four main types of stellar sources have been proposed for presolar grains: AGB (asymptotic giant branch) stars, RGB (red giant branch) stars, novae, and supernovae (Lodders & Amari, 2005). Nanodiamonds are the most abundant presolar grains, and they were the first isolated (Lewis et al., 1987). However, because they are also the smallest in size (~2 nanometers), limited by atom-counting statistics, their isotopic compositions are challenging to precisely analyze with current instrumentations (e.g., NanoSIMS). The stellar origins of nanodiamonds are thus not as well-studied as other types of presolar grains (e.g., SiC grains which are usually much bigger). Some (if not most) of the nanodiamonds may even be solar. Presolar nanodiamonds are possibly condensed from AGB or supernovae. Presolar silicates include olivine and pyroxene, and come mostly from RGB or AGB stars (86%), supernovae (10%), and others (Floss & Haenecour, 2016). Silicon carbides (SiC) are among the most well-studied presolar grains because of their large size (up to 20 micrometers) and the ease of separating them from the surrounding material. Most (93%) SiC grains (mainstream grains) are from low-mass (1.5–3 solar mass) AGB stars, and only 1% of all SiC grains (X grains) have strong isotope evidence of their supernovae origins (Amari, 2017). Oxides such as spinel, corundum, and hibonite can be categorized into four groups according to their O isotopic compositions. Different groups of presolar oxide grains can be explained by different mechanisms at different localities of RGB or AGB stars in various masses and metallicities (Zinner, 2014).

Future studies of presolar grains can focus on two directions (Amari, 2014): (1) searching for other types of presolar grains that were predicted by astrophysical models; or (2) analyzing isotopic compositions of elements heavier than Fe, which is difficult with current instruments due to the small sizes of presolar grains. Many different minerals are predicted by stellar nucleosynthesis and dust condensation models to condense from different stellar environments; more and more types of presolar grains have been discovered in primitive meteorites. A recent example is the identification of two unique presolar graphite grains: one containing an iron sulfide inclusion, representing the observation of a first iron sulfide presolar grain; the other graphite having isotopic compositions consistent with formation in the ejecta of a CO Nova binary star (Haenecour et al., 2016). New instruments have also been designed and constructed for high-precision measurements of heavy element isotopes in small samples like presolar grains. For example, the resonance ionization mass spectrometry (RIMS) has useful yield of 30–40% compared to 1% of NanoSIMS, which makes analyzing heavy element isotopes (e.g., Sr, Ba, Fe, and Ni) in high precision possible for presolar grains (Stephan et al., 2016).

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Figure 9. Secondary electron image of a graphite presolar grain separated from the Murchison (CM2) chondrite.

Courtesy of Pierre Haenecour.

Formation Age of the Solar System and Chronology of Meteorites

The basic principle of radiometric age dating was established in the first half of the 20th century. It is based on the constant rate of decay of radioactive isotopes regardless of their physical and chemical environment. Based on this principle, and the determination of decay constants, the age of the solar system (and also the age of the Earth) was first dated as 4.55 Ga by Clair Patterson in 1955 through analysis of the 207Pb/206Pb ratios of five iron and stony meteorites (Patterson, 1955; 1956). Although analytical precision has significantly improved over the past 50 years, this 4.55-Ga age has been virtually unchanged since then (cf. 4.56 Ga; Amelin, 2002). Through several decades’ efforts of dating the crystallization ages of different bulk meteorites and even various components within the same meteorites, we have constructed a timeline for the formation and evolution of the solar system.

The most commonly used extant radionuclide systems for dating meteorites are K-Ar (half-life: 1.27×109 years), Rb-Sr (4.88×1010 years), Sm-Nd (1.06×1011 years) and U-Th-Pb (238U: 4.47×109 years; 235U: 7.04×108 years; 232Th: 1.40×1010 years). For iron meteorites containing few silicates, the siderophile element Re-Os (4.16×1010 years) system is also particularly useful. All five of these systems rely on long-lived radionuclides (40K, 87Rb, 147Sm, 187Re, 238U, 235U, and 232Th) that still exist and are still currently decaying. The ages determined using these long-lived systems are called absolute ages. Several short-lived radionuclides are also often used for dating meteorites, such as 26Al (half-life: 7.3×105 years), 53Mn (3.7×106 years), and 182Hf (8.9×106 years). These isotopes existed at the beginning of the solar system; however, they have now completely decayed away to 26Mg, 53Cr, and 182W, respectively. Nevertheless, by measuring the concentrations of their decay products (e.g., 26Mg for 26Al), these short-lived radionuclides can be used to date the relative age differences of early solar system materials. As a result, ages calculated using these short-lived chronometers are called relative ages (however relative ages can be converted to absolute ages assuming one of the two materials being compared also has a calculated absolute age). All these isotope systems each yield different uncertainties in ages due to different half-lives, different relative precisions in the knowledge of the half-lives, different chemical and isotopic analytical precisions in the laboratory, and other factors (e.g., assumption of initially homogenous distribution for 26Al). The application of different isotope systems is also limited by the mineralogy and bulk chemistry of the meteorites (whether the meteorites/minerals contain enough or too much of certain isotopes). Currently the Pb-Pb double spike method has the best time resolution among long-lived radionuclide systems and can yield an age with an uncertainty of less than one million years (Amelin et al., 2009; Connelly et al., 2017). The short-lived radionuclide systems are also capable of producing relative ages (the time difference between two early solar system materials) with a precision of less than one million years (Davis & McKeegan, 2014). By combining different long and short-lived radionuclide systems, we can develop an accurate and precise timeline of the solar system formation (see Figure 10).

Calcium-aluminum-rich inclusions (CAIs) in carbonaceous chondrites are the oldest solids in the solar system. They contain extremely refractory minerals, which is consistent for the first condensates from the solar nebula. Their ages have been dated as 4567.30 ± 0.16 Ma using the most precise Pb-Pb dating method (Amelin, 2002; Amelin et al., 2010; Connelly et al., 2012; Connelly et al., 2017). This age defines the minimum age of the solar system. CAIs also have the highest inferred initial concentrations of the short-lived 26Al and 182Hf nuclides (Jacobsen et al., 2008; Burkhardt et al., 2008; Larsen et al., 2011), which is consistent with the long-lived systems and indicates that CAIs were formed before any other surviving solar-system solid phases. Other refractory inclusions, such as amoeboid olivine aggregates (AOAs), likely formed at the same time as CAIs based on their indistinguishably high 26Al/27Al ratios (Huss et al., 2001; Itoh et al., 2007). Chondrules, another major chondritic component, also began forming near simultaneously with CAIs. However unlike CAIs and AOAs, the formation of chondrules lasted for a few million years, likely due to reoccurring transient heating events such as nebular shock waves or impacts (Amelin, 2002; Connelly et al., 2012; Krot and Nagashima, 2017; Connelly et al., 2017; Budde et al., 2018). Previously, chondrules were thought to have formed 1–2 Ma after CAIs based on the 26Al/27Al difference between chondrules and CAIs (Kita et al., 2005); however, this 1–2 Ma-age gap has been eliminated by new measurements of chondrules made using the absolute age dating method (Connelly et al., 2012). However, unlike the Pb-Pb ages of 0–3 Ma after CAIs for chondrules across all groups (excluding CB chonrules), Al-Mg isotopic data provide narrow chondrules age ranges for different groups, for example, 2–3 Ma in CO, CV, and LL chondrites, 3.5 Ma in CR chondrites (Krot & Nagashima, 2017). This disagreement between the chondrules’ ages infered from the Al-Mg short-lived system and from the Pb-Pb long-lived system is still under debate and cannot be explained entirely by the heterogeneous distribution of initial 26Al abundances in the solar nebula (Budde et al., 2018).

Chondritic components such as CAIs, AOAs, and chondrules can be dated individually using different radionuclide systems suitable for their sizes, and their mineral and chemical compositions. These chondritic components, formed at different times and localities by various mechanisms in the solar system, eventually accreted in the region of the asteroid belt into chondritic parent bodies. The accretion ages of chondrites can only be inferred indirectly as they must be younger than the ages of individual chondritic components (CAIs, AOAs, and chondrules) and older than the timing of secondary processes (metamorphism or aqueous alteration) on the parent bodies. The timing of metamorphism (for ordinary chondrites) and aqueous alteration (for carbonaceous chondrites) on parent bodies can be dated via the minerals formed or affected (by resetting the isotopes) during metamorphism or aqueous alteration. These metamorphism or aqueous alteration ages on different chondritic parent bodies dated from secondary minerals vary from a few to tens of millions of years after CAIs (Keller et al., 1991; Endress et al., 1996; Hutcheon et al., 1998; Zinner & Göpel, 2002; Wadhwa et al., 2006; Hoppe et al., 2007; Hohenberg & Pravdivtseva, 2008; Trinquier et al., 2008; Nyquist et al., 2009; Bogard, 2011; Doyle et al., 2015).

Differentiated planetesimals such as the angrite parent body, the eucrite parent body, and some ungrouped achondrite (Asuka 881394) parent bodies accreted as early as 3 Ma after CAIs (e.g., 4564.4 ± 0.1 Ma, Amelin, 2008b, 2008a). Similarly, ages of iron meteorites (the cores of planetesimals) agree well with the early differentiation of angrite parent body and the eucrite parent body. “Non-carbonaceous” iron meteorites were formed at ~0.3–1.8 Ma after CAIs, while “carbonaceous” iron meteorites were formed at ~2.2–2.8 Ma based on 182Hf-182W ages (Kleine et al., 2009; Kruijer et al., 2017). The ages of accretion of planetary bodies such as Earth, Mars, and the Moon cannot be directly dated but can also be constrained by their core formation (the planetary differentiation) ages using the 182Hf-182W short-lived system (Kleine & Walker, 2017). The formation age of the martian core is ~10 Ma after CAIs (Dauphas & Pourmand, 2011). For the Earth, the formation age of the core is estimated to be >34 Ma after CAIs (Kleine & Walker, 2017). The final accretion of the Earth’s core is linked to the Moon-forming Giant Impact, where the core of the proto-Earth merged with that of the impactor (Canup & Asphaug, 2001). The age of the Moon-forming Giant Impact (and the age of the Moon) is still highly debated and model-dependent. The estimated ages range from 34 to ~100 Ma after CAIs (Lee et al., 1997; Halliday, 2000; Yin et al., 2002; Jacobsen, 2005; Kleine et al., 2009; Borg et al., 2011; Jacobson et al., 2014; Carlson et al., 2014; Bottke et al., 2015; Barboni et al., 2017; Kleine & Walker, 2017). The ages of Earth, Mars, and the Moon must be older than those of the rocks or minerals formed on them. The oldest preserved crust of the Earth is the detrital zircon from Jack Hills, Western Australia, which can be as early as ~185 Ma after CAIs (Valley et al., 2014). In contrast, the oldest preserved crust (lunar ferroan anorthosite) on the Moon was formed ~200 Ma after CAIs (Borg et al., 2011). On Mars, the oldest preserved crust (depleted shergottites) has an age of ~250 Ma after CAIs (Bouvier et al., 2009). In general, the timescales of planetary accretion and differentiation are extremely short compared to the 4.56-billion year history of the solar system.

Future studies of the chronology of the early solar system improve both long and short-lived radionuclide systems. The precision and accuracy of any radionuclide dating systems are affected by the following factors (Amelin, 2006): mass spectrometer sensitivity and stability; correction of instrumental mass fractionation; removal or subtraction of non-radiogenic daughter nuclides (including from nature or laboratory contamination); and disturbances of the closed-system by thermal and shock metamorphism. For any radionuclide systems, the uncertainties of decay constants affect the precision of calculated ages. For example, the uncertainty of the estimations of U isotopes decay constants cause a ~9 million-year error for Pb-Pb ages (Amelin, 2006). The uncertainty created by the decay constant does not affect the relative age between two samples that are both dated by the Pb-Pb method if they use the same decay constant. Occasionally, the decay constant is redetermined and significantly changed with new measurements. For example, the half-life of 146Sm was reanalyzed as 68 ± 7 million years instead of the previous value of 103 ± 5 million years (Kinoshita et al., 2012). The most precise long-lived radionuclide system is the stepwise-leaching double-spike U-isotope-corrected Pb-Pb method, which can yield an age with an uncertainty of ±0.25 Ma at best (Connelly et al., 2017). In comparison, other long-lived radionuclide systems, such as Rb-Sr and Sm-Nd, still do not have enough precision to be used to study the earliest events of the solar system. For short-lived radionuclide systems, the most widely used system is 26Al-26Mg. This dating system relies on the assumption that the 26Al abundances are homogenous in the solar nebula, which has recently been shown to be unlikely based on inconsistent Al-Mg and Pb-Pb ages (Bollard et al., 2017) and there is a gradient of 26Al from the inner to the outer solar disk (Larsen et al., 2016). Thus the application of short-lived radionuclide systems needs to be used with extreme caution and different radiometric dating methods need to be further cross-calibrated. However, whether the initial 26Al abundances are homogenous or not in the solar nebula is still debatable because a new study shows that the Al-Mg and Hf-W ages agree well with each other (Budde et al., 2018). It is still an ongoing problem that needs to be reconciled.

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Figure 10. Schematic timeline of the early solar system formation and planetesimal accretion (see main text for references).

The Building-Block Materials of the Earth

One of the key goals of meteorite studies is to identify the building blocks of the Earth and how they accreted to form the Earth. Since their formation at ~4.6 Ga, big planetary bodies like Earth have experienced global magmatic differentiation, which have erased all the initial structural and chemical information of the building-block materials. However, some extremely old and primitive meteorites such as chondrites, which formed early in the solar system, may have accreted from the same materials as those that formed the Earth. Even though the textual and chemical signatures of the Earth-building material have been changed, the mass-independent isotopic ratios (e.g., Δ‎17O, ε‎48Ca, ε‎50Ti, and ε‎54Cr) are still preserved. Isotopic anomalies are the remaining isotopic variation after the correction of mass-dependent fractionation, and they are caused by either nucleosynthesis or mass-independent fractionation processes. Isotopic anomalies were formed in the early solar system and they are not easily modified by subsequent chemical and physical processes on Earth and other planetary bodies. These isotopic anomalies can be used as the “fingerprint” or “DNA” to link genetic relationships between the Earth and meteorites.

Oxygen isotopes in meteorites are the first and most-widely studied isotopic systems that have been applied to trace the ancestry of different meteorites and planetary materials. The three isotopes of oxygen (16O, 17O, and 18O) can be fractionated mass-independently (Thiemens & Heidenreich, 1983; Clayton, 1993). If the solar system-building materials from many stellar sources (see the section “Presolar Grains in Meteorites and Stellar Sources of the Solar Nebula”) were sufficiently homogenized at the beginning of the solar system and if no mass-independent isotopic fractionation occurred in the solar nebula, all solar system materials should fall along the same mass-dependent fractionation line through one point representing the well-mixed initial composition of solar system. However, the measurements of meteorites and planetary materials show that they do not follow the same mass-dependent fractionation line (see Figure 5), and some of them even follow a mass-independent fractionation line (Clayton et al., 1973; Thiemens & Heidenreich, 1983). These observations indicate that the solar system-building materials from different stellar sources were not completely homogenized (Clayton et al., 1973) and/or there is mass-independent isotopic fractionation during photochemical processes (e.g., self-shielding of CO due to UV photo-dissociation) that occurred early in the solar nebula (Clayton, 2002; Yurimoto & Kuramoto, 2004; Lyons & Young, 2005). Different planetary bodies and asteroids have accreted materials from various regions in the solar system, thus have inherited different O isotopic compositions (see Figure 6). Besides oxygen isotopes, several other isotope systems such as Ca, Ti, Cr, Mo, Ni, Mo, Ru, and Nd also show isotopic variation (a.k.a. isotopic anomalies) after correcting all the mass-dependent fractionation (Dauphas & Schauble, 2016). Those Ca, Ti, Cr, Mo, Ni, Mo, Ru, and Nd isotopic anomalies are most likely caused by nucleosynthetic processes and carried by presolar grains.

These isotopic anomalies can be used to link different meteorites to the same parent bodies. As shown in the section “Historical Scientific Investigations of Meteorites,” oxygen isotopes are one of the most reliable criteria to classify meteorites. Some meteorites, such as primitive stony meteorites winonaites and IAB iron meteorites (Benedix et al., 2000) and undifferentiated meteorite enstatite chondrites and differentiated achondrite aubrites (Clayton and Mayeda, 1996) with different petrological and chemical features, can be categorized into the same group and be linked to the same asteroid. Other meteorites like main-group and Eagle Station pallasites, despite bearing similar mineralogy and geochemistry, are thought to be from two different parent bodies based on their different oxygen isotope compositions.

The isotopic composition of Earth is compared with various planetary materials to shed light on its initial building blocks. As shown in Figures 5 and 6, the isotopic compositions of Earth are identical to those of the Moon in all isotope systems showing isotopic anomalies, especially in ultra-high-precision oxygen isotopes (Wiechert et al., 2001; Young et al., 2016). This isotopic similarity between Earth and Moon is a curious issue that has fundamental implications for their origins. The currently most accepted theory for the origin of the Moon is the Giant Impact hypothesis (Hartmann & Davis, 1975; Cameron & Ward, 1976). The canonical Giant Impact scenario predicts that majority of the Moon comes from an impactor rather than from the Earth (Canup, 2004). The identical isotopic composition between the Earth and Moon requires either that the impactor have identical isotopic composition as the Earth (Dauphas, Burkhardt, Warren, & Fang-Zhen, 2014), or the canonical Giant Impact scenario needs to be revisited and revised (Canup, 2012; Ćuk & Stewart, 2012; Lock & Stewart, 2017). A new study of the isotopes of moderately volatile element K indicates that the Moon may have formed from a much more violent Giant Impact than the canonical Giant Impact scenario (Wang & Jacobsen, 2016; Lock et al., 2018). The new scenarios of Giant Impact and the origin of the Moon are currently at the frontier of the research of planetary formation.

Other than the Moon, enstatite chondrites and achondrites (aubrites) have nearly identical isotopic anomalies as Earth (see Figures 5 and 6). Earlier measurements of oxygen isotopes showed that enstatite chondrites and achondrites are indistinguishable from the Earth (Clayton et al., 1984b; Newton et al., 2000); however, recent high-precision measurements show a tiny but resolvable difference between enstatite chondrites/achondrites and Earth (Herwartz et al., 2014; Greenwood et al., 2018). Nevertheless, the oxygen isotopes of enstatite chondrites and Earth are incredibly close compared to any other meteorites. In addition, in all other isotopic anomalies (ε‎48Ca, ε‎50Ti, ε‎54Cr, ε‎64Ni, and ε‎92Mo) there is no measurable difference between Earth and enstatite chondrites (Dauphas, 2002; Trinquier et al., 2007; Regelous et al., 2008; Trinquier et al., 2009; Qin et al., 2010; Burkhardt et al., 2011; Zhang et al., 2012; Steele et al., 2012; Tang & Dauphas, 2012; Dauphas, Chen et al., 2014; Mougel et al., 2018). For ε‎100Ru and μ‎142Nd, enstatite chondrites are the closest to the Earth compared to other meteorites (Burkhardt et al., 2016; Fischer-Gödde & Kleine, 2017). This extreme and unique similarity between enstatite chondrites and Earth stimulated the proposal of the Enstatite Chondrite (EH) Model of the Earth (Javoy, 1995). Enstatite chondrites and the Earth likely accreted similar materials from the same region in the solar system where the isotopic composition of the materials was generally homogenous. Although the isotopic compositions of enstatite chondrites and the Earth are almost identical, their chemical compositions are different: (1) the major elements Mg/Si ratio of Earth are considerably higher than that of enstatite chondrites; (2) the Earth is depleted in volatile elements (e.g., K/U and Rb/Sr) compared to enstatite chondrites; (3) the Earth is much more oxidized than enstatite chondrites. Thus, although Earth and enstatite chondrites are nearly identical in all isotopic anomalies (Δ‎17O, ε‎48Ca, ε‎50Ti, ε‎54Cr, ε‎64Ni, ε‎92Mo, ε‎100Ru, and μ‎142Nd), and they probably share the same precursor from the early solar system, they clearly have experienced different histories of chemical evolution. In conclusion, Earth is not made of any known type of meteorites (but bears a close genetic relationship with enstatite chondrites). It is logical because possibly all Earth-building materials have already accreted into the Earth and none have left (at least in the current collection of meteorites).

Acknowledgments

Editor Bruce Fegley Jr. and one anonymous reviewer are thanked for their thorough reviews and comments. We appreciate help from Dr. Maxwell Thiemens, Dr. Pierre Haenecour and Dr. Piers Koefoed for their thorough proofreading and many suggestions.

Further Reading

Davis, A. M. (Ed.). (2014). Treatise on Geochemistry. Vol. 1: Meteorites and Cosmochemical Processes (2nd ed.). Oxford, U.K.: Elsevier.Find this resource:

Heide, F., & Wlotzka, F. (1995). Meteorites: Messengers from space. Berlin, Heidelberg: Springer.Find this resource:

Hutchison, R. (2007). Meteorites: A petrologic, chemical and isotopic synthesis. Cambridge, U.K.: Cambridge University Press.Find this resource:

Krinov, E. L. (1966). Giant meteorites. Oxford, U.K.: Pergamon Press.Find this resource:

Lodders, K., & Fegley, B., Jr. (2011). Chemistry of the solar system. London, U.K.: RSC Publishing.Find this resource:

McCall, G. J. H., Bowden, A. J., & Howarth, R. J. (Eds.). (2007). The history of meteoritics and key meteorite collections: Fireballs, falls and finds. Special publication. London, U.K.: Geological Society.Find this resource:

McSween, H. Y., & Huss, G. R. (2010). Cosmochemistry. Cambridge, U.K.: Cambridge University Press.Find this resource:

Papike, J. J. (Ed.). (1998). Reviews in mineralogy. Volume 36: Planetary materials. Washington, DC: Mineralogical Society of America.Find this resource:

Planetary Science Research Discoveries (PSRD).

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